High-level landscapes along the margin of southern East Greenland—A record of tectonic uplift and incision after breakup in the NE Atlantic

Global and Planetary Change 116 (2014) 10–29

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High-level landscapes along the margin of southern East Greenland—A record of tectonic uplift and incision after breakup in the NE Atlantic Johan M. Bonow a,b,⁎, Peter Japsen b, Troels F.D. Nielsen b a b

Södertörn University, SE-141 89 Huddinge, Sweden Geological Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, 1350 Copenhagen, Denmark

a r t i c l e

i n f o

Article history: Received 7 March 2013 Received in revised form 12 December 2013 Accepted 22 January 2014 Available online 29 January 2014 Keywords: East Greenland peneplain uplift erosion surface subsidence passive margin cenozoic Norway North Atlantic denudation chronology

a b s t r a c t Elevated plateaux and deeply incised valleys characterise the large-scale landscapes along the East Greenland margin as in many elevated, passive continental margins around the world. The absence of syn- or post-rift rocks in, for example, the mountains of Norway, hampers the assessment of the age of these landscapes and of the present-day elevation. The mountains of southern East Greenland (68–71°N), however, expose thick basalts that were extruded onto a largely horizontal lava plain near sea level during breakup of the NE Atlantic at the Paleocene–Eocene transition. We take advantage of these favourable geological conditions to investigate the uplift history after continental breakup. In particular, it is clear that present-day elevations of these basalts up to 3.7 km above sea level (a.s.l.) were reached after breakup. We have mapped regional erosion surfaces and integrated the information about the landscape with the stratigraphic record (i.e. stratigraphic landscape analysis). The analysis led to the following relative denudation chronology for southern East Greenland: At breakup, the margin subsided and underwent km-scale burial. Around the Eocene–Oligocene transition, the first phase of uplift, tilting and subsequent erosion led to the formation of an extensive, low-relief erosion surface (the Upper Planation Surface, UPS) that was graded towards the base level of the adjacent ocean before the eruption of Miocene lavas onto that surface. A second uplift that most likely occurred after the Miocene produced a new erosion surface (the Lower Planation Surface, LPS) by incision below the UPS. Finally, a third event in the late Cenozoic lifted the UPS and the LPS to their present elevations of up to 3 and 2 km a.s.l., respectively and shaped the present-day valleys and fjords by incision of rivers and glaciers below the LPS. The general picture of landscape development is highly similar to West Greenland and the common characteristics between the stepped landscapes in East Greenland and those on the conjugate margin in Scandinavia lead us to conclude that the mountains of Norway also formed after the North Atlantic breakup. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The large-scale landscapes in East Greenland are characterised by an elevated plateau at 2 to 3 km above sea level (a.s.l.) and by deep valleys incised below the main plateau (Figs. 1–4; Ahlmann, 1941; Brooks, 1979, 1985). Such landscapes with stepped surfaces are common along passive margins worldwide (Jessen, 1943; Japsen et al., 2012a; Green et al., accepted for publication), but the origin of such elevated plateaux remains a topic of great controversy, just like the origin of the elevated passive continental margins. Alternative viewpoints suggest that the elevated margins are: (a) remnants of old orogens (Nielsen et al., 2009a); (b) permanent uplifts somehow related to rifting and break-up (Swift et al., 2008; Sacek et al., 2012); and (c) young features caused by episodic burial and uplift after rifting and breakup and that their formation is due to build-up of stress related to changes in plate motion long after breakup (Japsen et al., 2012a,b).

⁎ Corresponding author. Södertörn University, Huddinge, SE-141 89, Sweden. E-mail addresses: [email protected], [email protected] (J.M. Bonow). 0921-8181/$ – see front matter © 2014 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.gloplacha.2014.01.010

The controversy can be exemplified with the debate about the origin of the margins around the North Atlantic with conflicting views on West Greenland (Redfield, 2010; Green et al., 2011), East Greenland (Thomson et al., 1999; Johnson and Gallagher, 2000; Swift et al., 2008; Pedersen et al., 2012; Japsen et al., 2013) and Scandinavia (Lidmar-Bergström and Bonow, 2009; Nielsen et al., 2009b; Chalmers et al., 2010; Gabrielsen et al., 2010; Steer et al., 2012; Hall et al., 2013). The key issue concerns the elevated topography with the plateau surfaces: when and how were the surfaces formed, when did they reach their present elevation and what does the answer to that question tell us about the properties of the Earth at depth? Whereas the age of rocks can readily be determined using radioactive methods, the age of mountains as topographic features cannot be easily estimated. This is particularly difficult in many areas around the North Atlantic compared to other continental margins because glacial erosion has often removed any cover (if ever deposited) that could have defined the maximum age of the topography. In contrast, the study area in southern East Greenland, between Kangerlussuaq and Scoresby Sund (c. 68–71°N; Figs. 2–4), contains Jurassic–Neogene

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Fig. 1. Elevation of bedrock onshore and bathymetry offshore in the North Atlantic domain. A common feature of the elevated margins is that they have one or several elevated plateau surfaces that are dissected by deeply incised valleys; e.g. Norway (Gjessing, 1967; Lidmar-Bergström et al., 2000, 2013), Scotland (Hall, 1987), West Greenland (Bonow et al., 2006a,b), Baffin Island (Kleman, 2008) and East Greenland (Ahlmann, 1941; this study). Note the much higher elevation in East Greenland compared to West Greenland. The load of the Greenland ice sheet causes a depression of up to 800 m in central Greenland whereas peripheral bulging caused by the ice load has a negligible effect on the elevation of Greenland's margins (Medvedev et al., 2013). Study area in southern East Greenland marked by dashed line. Elevation data source: Amante and Eakins (2009).

rocks which document the conditions prior to, during and after the onset of spreading in the NE Atlantic at the Paleocene–Eocene transition (~ 56 Ma), when massive flood basalts were extruded across the area during rapid subsidence (Fig. 5; Brooks and Nielsen, 1982; Nunns, 1983; Larsen et al., 1989; Pedersen et al., 1997; Larsen and Saunders, 1998; Larsen and Tegner, 2006; Henriksen et al., 2009; Brooks, 2011). Today, marine mid-Paleocene sediments of the Sediment Bjerge Formation occur at elevations up to 1.4 km a.s.l. near Pyramiden in the Kangerlussuaq Basin (A. Whitham, pers.comm.; 2013). Southern East Greenland is thus highly suitable for deciphering the tectonic and landscape history. This study takes advantage of these favourable geological conditions to investigate the uplift history after continental breakup. We do so by identifying and mapping erosion surfaces and by integrating the information about the landscape with the stratigraphic record; i.e. stratigraphic landscape analysis as introduced by Lidmar-Bergström et al. (2013). This approach enables us to put tight constraints on a relative denudation chronology for central East Greenland that defines when the plateau surfaces were formed and when they reached their present elevation. In a paper parallel to this (Japsen et al., 2014–in this issue), the results presented here are integrated with thermochronological data. We finally discuss the implications of our results in this study for other highly elevated passive continental margins that do not have the same detailed geological control.

2. Observations of large-scale landscapes as input for conclusions about base-level changes in the past Extensive upland plains around the globe have been interpreted in classic geomorphological papers as erosional surfaces graded towards the ultimate base level and subsequently uplifted (Davis, 1899; Penck, 1924; King, 1967). Similar observations and conclusions were made by geomorphologists during the 20th century for the margins in the North Atlantic domain, including Scandinavia, Scotland, Newfoundland, Baffin Island and East Greenland (Fig. 1; Reusch, 1901; Ahlmann, 1919, 1941; Jessen, 1943; Holtedahl, 1953; George, 1966; Brookes, 1977). They all regarded the elevation of the plateaux and the valleys incised below the plateaux to be the result of uplift during the late Cenozoic. Similar conclusions were reached in later studies of large-scale landforms in West Greenland (Bonow et al., 2006a,b), in Scandinavia (Lidmar-Bergström et al., 2007, 2013) and NE Brazil (Bonow et al., 2009). More recently, Steer et al. (2012) suggested that low-relief surfaces at high elevation in western Norway have been formed by glacial head-ward erosion by cirque retreat, although this was refuted by Hall et al. (2013), who showed evidence that glacial erosion acts to dissect plateaux rather than create them. We have used detailed mapping of extensive erosion surfaces that cut across rocks of different age and resistance to investigate the evolution of elevated passive continental margins. Analysis of landscapes in Scandinavia (e.g. Lidmar-Bergström et al., 2000, 2007, 2013; Bonow

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Fig. 2. Geological map of southern East Greenland. Modified after Larsen et al. (1989, 2002, 2005a,b), Pedersen et al. (1997), Tegner et al. (1998, 2008), Storey et al. (2004) and Henriksen et al. (2009). Locations of photos in Fig. 3 and profile in Fig. 10 are indicated. PWB: Prinsen af Wales Bjerge.

et al., 2003), West Greenland (Bonow, 2005; Bonow et al., 2006a,b) and NE Brazil (Bonow et al., 2009; Japsen et al., 2012b) has led us to conclude that fluvial erosion to base level is fundamental for understanding the formation of extensive, low-relief surfaces. Base levels control fluvial erosion and the ultimate base level is sea level. A local base level can also be highly important as, for example, a resistant layer or in an internal drainage basin (Fjellanger and Etzelmüller, 2003; Babault et al., 2005). Where a study area is known to have been near the sea at the time a peneplain formed (as in the case of the post-rift development of margins adjacent to opening oceans) the most obvious choice of base level is sea level (see Japsen et al., 2009). Resistant rocks (as described above) will affect the transient landscape, but their influence will diminish through time (e.g. Fig. 2 in Japsen et al., 2009). The process of valley widening and the removal of material by river systems (e.g. Ahnert, 1998; Bonow et al., 2007; Lidmar-Bergström et al., 2007) eventually results in a large-scale, low-relief erosion surface with a low slope gradient; a planation surface (one type of peneplain). We describe such denudational planes of low relative relief (b 200 m), either horizontal or inclined, with different characteristics as flat, hilly, or flat with scattered hills as peneplains (see Green et al., 2013). Valley incision below a peneplain surface is evidence of further lowering of base level (uplift of the landmass or drop in sea level), with subsequent formation of new valley floors and valley widening, grading to sea level and thus possibly to the formation of a new low-relief erosion surface. Landscapes characterised by surfaces that appear in steps are common on all continents, not only in formerly glaciated areas (Japsen et al., 2009, 2012a; Green et al., 2013). The height difference between the valley floor and

the overlying surface therefore indicates the amount of uplift or drop in base level (Fig. 6). We apply stratigraphic landscape analysis which is based on: (a) the relationship between peneplains in crystalline basement and their cover rocks of different ages; (b) the cross-cutting relationships between such re-exposed peneplains and epigene peneplains (ones that have never been covered); and (c) the occurrence of valleys incised below peneplains (see Fig. 7). Where possible, e.g. in East Greenland, we also study peneplains across sedimentary or volcanic rocks. Stratigraphic landscape analysis has been successfully developed and used in Scandinavia e.g. by Lidmar-Bergström et al. (2013) to identify different Phanerozoic episodes of erosion, uplift and subsidence (cf. LidmarBergström, 1982, 1988, 1996; Lidmar-Bergström et al., 2000, 2007). Japsen et al. (2006) provided new insight into the nature of uplifted peneplains in a study from central West Greenland that integrated observations from the geological record with the results of stratigraphic landscape analysis (Bonow et al., 2006a,b) and thermal history reconstructions based on thermochronological data (Japsen et al., 2005). It was possible to demonstrate that the extensive planation surfaces in West Greenland formed after the start of sea-floor spreading in the Labrador Sea in the Paleocene that they were graded to base level (the adjacent sea) during the Oligocene–Miocene and that km-thick successions of rock were removed during peneplain formation. Consequently, it was concluded that the present-day elevation of these surfaces is due to uplift that happened after their formation and that uplift began in the late Miocene, many tens of millions of years after continental breakup.

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A

B

Fig. 3. Oblique aerial photographs illustrating the main characteristics of the landscape in southern East Greenland with large upland plains and deeply incised valleys. A: Gåseland. Looking north-west across the inner part of Gåsefjord to the extensive plateau (the Upper Planation Surface, UPS; Fig. 5) that here cuts across the Main Basalts (Geikie Plateau Formation, c. 1.9 km a.s.l.). The undulating basement surface (hilly relief) in the foreground is covered by basalts of the Milne Land Formation. The extent of elevated valley benches (the Lower Planation Surface, LPS) is restricted and can be identified only in the far background (cf. Fig. 11). Detail from oblique aerial photograph 654-G-N, no. 12159, 1950. © KMS, Denmark. B: Geikie Plateau. Looking north-west across a landscape characterised by sharp ridges and scattered, partly ice-covered remnants of the UPS that cuts across the Main Basalts (Skrænterne Formation) at the southern edge of the Geikie Plateau (c. 2 km a.s.l.). Location of photographs in Fig. 2. Detail from oblique aerial photograph B-NV 661–10649, 1948. © KMS, Denmark.

The observation that the large-scale landscapes of the West Greenland margin are similar in many respects to those of other elevated, passive continental margins around the world (Japsen et al., 2012a), leads us to the general suggestion that the formation of all such elevated margins is unconnected to rifting and breakup and that their topography formed later (Bonow et al., 2007; Japsen et al., 2012a). This view is contrary to earlier ideas that elevated passive margins are steadystate landscapes (e.g. Ollier and Pain, 1997; Brown et al., 2000; Bishop, 2007; Nielsen et al., 2009a,b). 3. Methods The study area was visited twice as part of the present investigation. In 2008 fieldwork was made on Milne Land, including helicopter reconnaissance further north and across Jameson Land. A field camp was established in the Kangerlussuaq area (Sødalen) in 2009 and helicopter flights were made with ground stops along Kangerlussuaq, into the area south of Kangerlussuaq, to Kap Edvard Holm and to Watkins Bjerge and Gunbjørn Fjeld in the north (Fig. 2).

We used a digital elevation model with c. 30-metre resolution (ASTER data) as input data (Fig. 4) and found that a 100 m contour map gives a reasonable picture of the general landscape features such as flats, escarpments and deeply incised valleys (cf. Bonow, 2004, p. 8–9). We constructed a contour map from the elevation data as the primary input for the surface mapping. A contour map shows the full threedimensional picture of the landscape whereas the relief along a topographical profile depends on the location and azimuth of the transect. To support the mapping we used profiles and we therefore divided the study area into a square grid of profiles, spaced 25 km apart. The topographical profiles extracted along the grid lines, as well as the contour map were plotted in a scale of 1:250,000, so that a direct comparison could be made during mapping. We also extracted maximum and minimum heights along these profiles in a swath 50 km wide. The swath profiles were plotted together with the topographical profile (Fig. 8). In total, we cross-analysed 32 profiles with a total length of c. 10,000 km. The mapping of the surfaces was initiated in core areas with low, relative relief and only minor valleys that were defined from the contour

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Fig. 4. 3D elevation model of southern East Greenland based on ASTER digital elevation data (c. 30 by 30 m resolution). The model is vertically shaded to enhance the slopes. The Wager Plateau and the Geikie Plateau, which combine to form a slightly tilted upland plateau with elevations between c. 2500 m a.sl. (purple) and 2000 m a.s.l. (red) are dominant features of the landscape in the central part of the study area and in the Scoresby Sund area. The plateaux are dissected by deeply-incised valleys which are filled with large glaciers at present. The length of the foreground of the figure is about 500 km and the geographical location of the 3D model can be seen from the index map and from Fig. 2.

S Kangerlussuaq

Scoresby Sund

Inland succession

N Milne Land Sediment

Coastal succession

Kap Brewster Fm ?Miocene, ?23–?5 Ma

LPS Vindtop Fm 14 Ma

Kap Dalton Group Lutetian, 47-44 Ma (Kap Dalton) ?Lutetian – Rupelian, ?48-28 (Savoia Halvø)

M

UPS BT 47 Ma Undated pictrites

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Kangerlussuaq Group Albian – Selandian,105 – ~60 Ma

Igtertivâ Fm 49 Ma M

KI 51 Ma

Charcot Bugt Fm Middle Jurassic, 170–160 Ma

PWB Fm 53 Ma

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Rømer Fjord Fm SI 56 Ma

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Late volcanics Ypressian and younger, <54 Ma

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Plume

Lower Basalts/Nansen Fjord Fm Selandian – Thanetian, ~60–56 Ma

M

Crystaline basement

Precambrian

Caledonian

Angular unconformity Peneplain Etch surface

Fig. 5. Stratigraphy of the main geological units as mentioned in the text and extent of the erosion surfaces (this study) for the Kangerlussuaq, Scoresby Sund and Milne Land area (68–70.5°N). Volcanic stratigraphy after Hansen et al. (2002) (based on Nielsen et al., 1981; Larsen et al., 1989, 1999, 2013; Hansen et al., 1995; Storey et al., 1996; Pedersen et al., 1997; Tegner et al., 1998; Heister et al., 2001; Storey et al., 2004). Sediment stratigraphy after Hassan (1953), Nunns (1983), Larsen et al. (2002; 2005), Nøhr-Hansen and Piasecki (2002), Surlyk (2003) and Nøhr-Hansen (2012). Tectonic events after Nunns (1983), Dam et al. (1998) and Nielsen et al. (2006). Ages of intrusive complexes after Riishuus et al. (2008) and Tegner et al. (2008). Archaean gneiss basement near Kangerlussuaq was reworked in the Proterozoic whereas the basement rocks around Scoresby Sund were reworked during the Caledonian orogeny (Higgins and Leslie, 2008; Nutman et al., 2008). Abbreviations. BT: Borgtinderne; KI: Kangerlussuaq Intrusion; LPS: Lower Planation Surface; PWB: Prinsen af Wales Bjerge; SI: Skaergaard Intrusion; UPS: Upper Planation Surface; LPS: Lower Planation Surface.

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Sea level uplift c. 2 km

uplift c. 1 km

Time Basement

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LPS

Fig. 6. Conceptual sketch illustrating the formation of the elevated plateaux in the study area. A: Extrusion of basalt during subsidence. B: Formation of an erosion surface (in this case, the Upper Planation Surface, UPS) by erosion to base level (the level of the adjacent sea) after uplift and block tilting; extrusion of the middle Miocene lavas of the Vindtop Formation (Storey et al., 2004) onto the UPS. C: Uplift of the UPS, with subsequent incision of valleys. D: Development of a new erosion surface (the Lower Planation Surface, LPS) by incision and by valley widening towards the newly formed base level. The UPS becomes dissected by fluvial backward erosion as well as by valley widening. E: Uplift of both UPS and LPS with subsequent incision of valleys below the LPS that later became reshaped by glacial erosion.

map (and with support from oblique aerial photographs where available). In such areas the maximum height along each swath coincides with the topographical profile. The same surface can also be mapped in areas that are much more dissected by glacial or fluvial erosion with support from the profiles. The edge of a surface is located where a rapid change of inclination is observed on the contour map. The exact slope angle used to define the rapid change varies due to bedrock strength, but in a study of a crystalline basement area in southern Norway, the slope was found to be 6.5° (Bonow et al., 2003) and a similar value was observed in a study in central West Greenland (Bonow et al., 2006b). The combination of profiles and contour map identifies offsets within a surface, e.g. across faults (Lidmar-Bergström, 1988; Bonow et al., 2006b). Surfaces incised below a plateau can be identified using the minimum height along profiles as such surfaces are controlled by fluvial erosion to a lower base level. In summary, we identified levels of the low-relief surfaces on the map and on the profiles. The interpretations made on the contour map were cross-checked from profiles to ascertain that the interpretations were consistent. Finally, mapped surfaces were compared with geological maps (Bengaard and Henriksen, 1984; Myers et al., 1988; GEUS, 2007) to assure that erosion surfaces could be differentiated from any structural surfaces.

4. Geological context To understand the large-scale landscape development in southern East Greenland, it is important to appreciate the tectonic history

Milne Land/Gåseland UPS

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4.1. Basement and cover rocks Precambrian and Caledonian metamorphic and intrusive rocks make up the crystalline basement around Kangerlussuaq and Scoresby Sund, respectively (e.g. Higgins and Leslie, 2008; Nutman et al., 2008; Henriksen et al., 2009). In the north, the study area overlaps with the post-Caledonian sedimentary basin in NE Greenland; a Triassic–Cretaceous succession crops out on southern Jameson Land and southern Liverpool Land and Middle Jurassic sandstones onlap the crystalline basement on eastern Milne Land (Larsen et al., 1989; Surlyk, 2003; Henriksen et al., 2009). In the Kangerlussuaq Basin, the Lower Cretaceous succession rests on the basement south of Christian IV Fault (below Christian IV Gletscher) whereas the Paleogene succession rests on the basement north of the fault (Larsen and Whitham, 2005). Throughout the area, the earliest basalt flows (see below) rest partly on the basement. The basement is deeply weathered to kaolinitic saprolites at the contact with the sedimentary cover in the Milne Land–Gåseland area (Birkelund and Perch-Nielsen, 1976; Larsen et al., 1989). When the saprolites are stripped from such a weathered surface, the surface is characterised by distinct hills separated by small valleys following the fracture systems. This hilly relief basement surface was buried below the early lavas and the basalts eventually also covered the basement highland on Milne Land up to 1800 m above the present sea level (Fig. 9). A few steep-sided valleys of pre-basalt age have been preserved on the border of a fault-controlled escarpment between the basement high and the Mesozoic basin (maximum relief 700 m; Larsen et al., 1989). 4.2. Sediments and volcanic rocks of the Kangerlussuaq Basin

LPS

Incised valley

revealed by the geological record and the post-volcanic structures (Figs. 2 and 5).

Jurassic

Cover

Fig. 7. Sketch of erosion surfaces and incised valleys and their relationships to basement and cover rocks in the Milne Land and Gåseland area. ES: etch surface (pre-Jurassic or pre-basalt age), UPS: Upper Planation Surface (post-basalt age), LPS: Lower Planation Surface (post-basalt and also post-UPS age).

North-east of Kangerlussuaq, the Kangerlussuaq Basin exposes a c. 1 km thick, Lower Cretaceous to Paleocene sedimentary and volcanic succession that accumulated on the Precambrian basement at the western margin of the seaway between Greenland and the British Isles (e.g. Wager, 1947; Soper et al., 1976; Nielsen, 1981; Nielsen et al., 1981; Larsen and Whitham, 2005; Larsen et al., 2005a,b; Nøhr-Hansen, 2012). The succession documents a transgression during an Albian phase of continental rifting, followed by shallow-marine deposition that dominated during the Cretaceous until intense rifting and deepmarine deposition occurred in the latest Cretaceous and early Paleocene (Larsen and Whitham, 2005; Nøhr-Hansen, 2012). Dam et al. (1998) argued that the impact of the Iceland plume controlled rapid uplift and fluvial erosion in the mid-Paleocene and that subsequent subsidence was contemporaneous with the first volcanic extrusions. The uplift caused rotation and erosional truncation of the sediments below

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Topography along

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Minimum topography along swath between coordinates 7610000 and 7660000

Fig. 8. Illustration of how low-relief surfaces can be identified from elevation data along swath profiles. A. Elevation contours in the Kangerlussuaq area and the area covered by the swath profile in B (grey rectangle). B. W–E profile showing the topography along UTM Y-coordinate 7635000 plus minimum and maximum topography within a 50 km wide swath around this profile (grey zone on map). Our interpretation of the LPS is indicated along the profile (see Fig. 11). Remnants of the UPS are preserved north of the swath (at Lindbergh Fjelde), but the summit of the Domkirkefjeldet is near the level of the UPS according to our interpretation. Contour interval 100 m. UTM coordinates, zone 24 northern hemisphere.

the earliest flows of the overlying Lower Basalts comprising plume-type picrites (Nielsen et al., 1981, 2006). The Selandian–Thanetian Lower Basalts are a formation of picritic to basaltic lavas and sediments that accumulated in the continental rift prior to the eruption of the latest Thanetian–earliest Ypressian Main Basalts (flood basalts) (Wager and Deer, 1939; Nielsen et al., 1981). The transition from pre-volcanic sedimentation to volcanic eruptions is well

documented in the Kangerlussuaq Basin. Nielsen et al. (1981) reported marine bivalves in sediments that were deposited immediately prior to extrusions of thick hyaloclastites (Mikis Formation, upper part of the Lower Basalts) and thus documented that the first hyaloclastites were partly extruded in a marine environment. The basal part of the Lower Basalts (Vandfaldsdalen Formation) also accumulated partly in a marine environment (Nøhr-Hansen, 2012); see Larsen et al. (2001), their Fig. 5.

NE

SW

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Ice surface

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Geikie Plateau Fm

Milne Land Fm

Rømer Fjord Fm

Geikie Plateau Skrænterne Fm v

v

Top of lava pile predicted from elevation of zeolite zones

Cover

Fig. 9. Profile showing the inclined and truncated strata of the Main Basalts and the original top of the lava pile estimated from the elevation of zeolite zones (Larsen et al., 1989) across Milne Land, Gåseland and Geikie Plateau (Table 1). The present-day lava surface is consequently an erosional feature and the tilt and offset of the basalt formations are post-basalt features. Based on Larsen et al. (1989). Location of profile in Fig. 13.

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4.3. Main Basalts

17

~2 km a.s.l., north of the head of Kangerlussuaq where both formations are capped by a fluviatile conglomerate and by lava flows of unknown age. Igtertivâ Formation, early Eocene, c. 49 and 44 Ma (Soper and Costa, 1975; Larsen et al., 1989, 2013; Tegner et al., 2008). This formation occurs in small grabens at Kap Dalton below and interbedded with partly marine sediments of the Bopladsdalen Formation of the Kap Dalton Group (see below). Sediment layers are commonly present between the lava flows and some layers contain marine microfossils (Fig. 10). The main part of the succession below the sediments of the Bopladsdalen Formation is dated at 49.09 ± 0.48 Ma while a lava flow intercalated with sediments of the Bopladsdalen Formation has an age of 43.77 ± 1.08 Ma (Larsen et al., 2013). The Main Basalts bounding the graben to the NW are nearly flatlying and belong to the Geikie Plateau, Rømer Fjord and Skrænterne Formations, indicating a downthrow of the graben succession in excess of 1500 m (Larsen et al., 1989). The lavas of the Igtertivâ Formation are not cut by any dykes, in contrast to the surrounding lavas where cross-cutting NE-running dykes are frequent (Watt, 1975). The majority of the dykes were intruded before the main graben-forming faulting took place (Watt et al., 1976) and because of compositional similarities these dykes were considered to have fed the Igtertivâ Formation (Larsen et al., 1989). Larsen et al. (2013) concluded that the majority of the faults that cut the succession at Kap Dalton and the formation of the exposed small graben are younger than the sediments preserved in the graben. The Main Basalts at Kap Dalton were thus below sea level at the time of the extrusion of the Igtertivâ Formation and the present elevation of the Main Basalts at this location consequently shows that rock uplift of at least 1500 m affected the margin after breakup as also pointed out by Clift et al. (1998).

The Main Basalts (flood basalts) erupted during the main volcanic phase that lasted less than 1 Myr and accompanied the breakup at the Paleocene–Eocene transition (~56 Ma; Larsen and Tegner, 2006; Storey et al., 2007). Net subsidence dominated the area during the eruption of the Main Basalts with no evidence for crustal upwarping at this time (Wager, 1947; Brooks and Nielsen, 1982; Brooks, 1985, 2011; Larsen et al., 1989; Pedersen et al., 1997; Larsen and Tegner, 2006). The succession attains a vertical thickness of more than 6 km in the central coastal area and thins inland to 2–3 km; the basalts cover about 65,000 km2 and reach far inland from Blosseville Kyst (Pedersen et al., 1997). Pedersen et al. (1997) investigated the lavas between Kangerlussuaq and Scoresby Sund by multi-model photogrammetry and profile sampling and established a three-dimensional framework for the lava geometry. They found no evidence for major unconformities within the Main Basalts whereas tilting of large blocks occurred far inland during the eruption of the Lower Basalts. Pedersen et al. (1997) observed syn-volcanic, small-scale faulting of only a few tens of metres within the Main Basalts and they noted that individual lava flows covered up to thousands of square kilometres (Larsen et al., 1989). Pedersen et al. (1997, p. 567) therefore concluded that the lava flows of the Main Basalts erupted onto a “largely horizontal lava plain without significant relief”. The four formations of the Main Basalts extend over much of the study area: Milne Land, Geikie Plateau, Rømer Fjord and Skrænterne Formations (Larsen et al., 1989; Pedersen et al., 1997). Pedersen et al. (1997) estimated that a differential sagging (N 2 km) of the coastal areas took place during the emplacement of the lavas and suggested that this indicated focusing of the magmatic production into a developing rift zone close to the present coast. A major coastal flexure developed in the Kangerlussuaq region by the collapse of the attenuated crust subsequent to the extrusion of the flood basalts (Nielsen and Brooks, 1981). 4.4. Younger lava formations Three basalt formations of limited extent post-date the Main Basalts (Fig. 5):

Vindtop Formation, middle Miocene, c. 14–13 Ma (Storey et al., 2004). These lava flows crop out on nunataks within an area of about 18 km2 between 2.7 and 2.9 km a.s.l.

Prinsen af Wales Bjerge Formation, early Eocene, c. 53 Ma (Hansen et al., 2002). These lavas overlie the Milne Land Formation at

Steno Bræ

25 km

M Marine sediment Mid-Miocene to Recent sediment Late Oligocene – mid-Miocene sediment Eocene – Late Oligocene sediment Subaerially erupted Paleogene basalt Oceanic basement Mesozoic sediment

Kap Dalton

3 2 1

Mid -Mio

M 5 4

3

M 5: Kap Dalton Group 4: Igtertivâ Fm 3: Skrænterne Fm 2: Rømer Fjord Fm 1: Geikie Plateau Fm

COT

cene

Unconfo rmity

3 2 1 0 1 2 3 4 5 6 7 8 9

Elevation (km)

E

Depth (km)

W

3 Maximum elevation in corridor UPS 2 1 Elevation along profile 0

Fig. 10. Onshore–offshore profile illustrating post-breakup, differential vertical movements of the Paleogene basalts. Studies of the Main Basalts along Blosseville Kyst show that they erupted onto a largely horizontal lava plain (Pedersen et al., 1997). The presence of marine incursions in some of the Lower Basalts (e.g. Nielsen et al., 1981) and in the uppermost of the flood basalts, the early Eocene Igtertivâ Formation (Soper and Costa, 1975; Tegner et al., 1998) also implies that the landscape was low-lying near sea level during the volcanic eruptions. Today the Igtertivâ Formation with marine incursions is exposed at Kap Dalton within a graben that was downfaulted from a position above the adjacent mountains where the lavas of the older lavas of the Skrænterne Formation are now exposed at 1500 m a.s.l. (Larsen et al., 1989; Larsen and Saunders, 1998). The Main Basalts at Kap Dalton were thus below sea level at the time of the extrusion of the Igtertivâ Formation. Drilling at ODP Site 118 documented subaerially erupted lavas 3.1 km below the seabed (~63°N; Larsen et al., 1993) and Larsen and Saunders (1998) document that the lavas off Blosseville Kyst also erupted subaerially. Consequently, the large lateral variation of the elevation/depth of the basalts along Blosseville Kyst represents differential, vertical movements after breakup. Offshore profile from Larsen (1990). Location in Fig. 2.

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J.M. Bonow et al. / Global and Planetary Change 116 (2014) 10–29

4.5. Intrusions

4.7. Elevation of the land surface during the eruption of the flood basalts

Tegner et al. (1998, 2008) compiled evidence about the timing of the mainly Eocene, alkaline intrusions along southern East Greenland that occur up to 100 km inland and found that they mainly crop out in tectonic and magmatic lineaments orthogonal to the rifted margin. South of Kangerlussuaq (c. 67–67.5°N), the intrusions are confined to three time windows at 56–54, 50–47 and 37–35 Ma. Along Kangerlussuaq (c. 68–68.5°N), the ages of the plutons span from 56 to 40 Ma and to the north (c. 68.5–69°N), intrusions range from 52 to 36 Ma. Two important intrusives are:

The Lower Basalts include basal hyaloclastites with marine incursions and waterlain tuffs that occur up to 2 km above the base of these flows and they are overlain by up to 5 km of flood basalts (Main Basalts) that are capped by marine sediments (Igtertivâ Formation and Kap Dalton Group) (see above). These observations demonstrate that the lavas accumulated near sea level during the earliest volcanic eruptions and that subsidence during the eruption of the Main Basalts kept pace with the outpouring lavas, so that the land surface at the end of the eruptions remained at or near sea level. Because the Main Basalts erupted onto a largely horizontal lava plain without significant relief (Pedersen et al., 1997), the conditions at Kap Dalton during the later eruption of the lava flows of the Igtertivâ Formation, can be considered as representative for Blosseville Kyst. The balance between subsidence and accumulation follows from the easy access from magma chambers in the crust to the eruption sites for the flood basalts that occur throughout the region. Tectonic reconstructions demonstrate that the crust was thinned at breakup (Nielsen and Brooks, 1981) and the emplacement of dyke swarms recalls the sheeted dyke complexes of ophiolites. The tectonic regime at that time was thus extensional and this resulted in subsidence of the base of the basalts. The flood basalts in East Greenland are all equilibrated in and tapped from replenishment, tapping and fractionation (RTF) magma chambers at crustal pressures (Larsen et al., 1989). The bulk compositions of the basalts are evolved and they have complementary cumulates within or at the base of the continental crust. As much as 50% of the original mass of mantle-equilibrated basalt would have crystallized to reach the composition of the extruded basalts. So, at the time when the lavas build up on the surface, they were tapped from magma chambers below the lava fields and neither rise nor subsidence of the land surface should be anticipated. At present, it is not possible to make up a total mass balance, but the stratigraphic evidence demonstrates that subsidence and accumulation of extrusives kept pace, so that the land surface remained near sea level during the eruption of the flood basalts. This conclusion has been reached by previous workers and for example Brooks (1985, 2011) concluded that the presence of marine fossils near the top of the lava succession in the Scoresby Sund region (Igtertivâ Formation and Kap Dalton Group) showed that the thickening of the lava pile was balanced by the deepening of the basin (Fig. 10). These results closely match the interpretation of the vertical motions of the spreading system offshore SE Greenland presented by Hopper et al. (2003). According to these authors, the ridge system was initially close to sea level for at least 1 Myr despite a reduction in magmatic productivity over this time interval. They explain this observation by dynamic support to the ridge by a small component of active upwelling into a pre-existing lithospheric thin spot as a thin sheet. Exhaustion of the thin sheet led to rapid subsidence of the spreading system and a change from subaerial to shallow marine and, finally, to deep marine extrusion in ~2 Myr.

Skaergaard Intrusion (55.6 Ma; Brooks, 2011). The intrusion is emplaced into the basal portion of the flood basalts (the Milne Land Formation) on the east side of Kangerlussuaq, near the present-day coast (e.g. Brooks and Nielsen, 1982; Larsen and Tegner, 2006). Larsen and Tegner (2006) investigated the pressure under which the intrusion crystallized–and thus its burial history–by analysing e.g. fluid inclusions in quartz and the mineral assemblage in the granophyres. They inferred that an estimated pressure increase from the basal to the upper part of the intrusion was due to progressive burial during the cooling of the intrusion and contemporaneous outpouring of 5–6 km of overlying flood basalts. Kangerlussuaq Alkaline Complex (50 Ma; Brooks, 2011). The main constituent is the Kangerlussuaq Intrusion which covers about 800 km2 on the west side of Kangerlussuaq (e.g. Wager, 1965; Riishuus et al., 2008). Riishuus et al. (2008) presented a model of the emplacement of the Kangerlussuaq Intrusion below a roof of 3 to 4 km of flood basalts.

4.6. Post-volcanic sediments Scattered remnants of a previously more widespread, fluvial to shallow-marine, post-volcanic succession are exposed at Kap Dalton and at Savoia Halvø (e.g. Hassan, 1953; Birkenmajer, 1972; Larsen et al., 1989; Birkenmajer and Jednorowska, 1997; Larsen et al., 2002, 2005a; Nøhr-Hansen, 2012). Kap Dalton Group; Bopladsdalen and Krabbedalen Formations (middle Eocene–early Oligocene; Larsen et al., 2002, 2005a). These fine-grained deposits are exposed at both locations and at Kap Dalton they rest on and inter-finger with the down-faulted lavas of the Igtertivâ Formation. Larsen et al. (2005a) studied the sediments at Kap Dalton and showed that the area suffered subaerial erosion subsequent to the extrusion of the basalts, leading to formation of an irregular relief, dissected by fluvial channels that eventually were buried below shallow-marine sediments during early middle Eocene transgression. The presence of clean quartz sandstones within the early post-basaltic succession indicates that areas with little or no basaltic cover became periodically exhumed. At Savoia Halvø, the marine sediments of the Krabbedalen Formation are of early Oligocene age (Larsen et al., 2002, S. Piasecki, pers. comm., 2009). Kap Brewster Formation (?Miocene; Hassan, 1953). These marine deposits overlie the basalts at one small locality on Savoia Halvø. The deposits consist of breccia, conglomerate and sandstone of dominantly basaltic material and are older than some of the faulting in the area (Larsen et al., 1989). Hassan (1953) tentatively assigned a Miocene age to these deposits based on the macrofossils and this age has still to be confirmed (Larsen et al., 2002).

4.8. Regional zeolite zones Zeolites are late stage minerals that fill cavities in basalts and the progress of zeolite minerals depend on temperature. Zeolite isograds delineating regionally extensive mineral zones occur throughout most of the Main Basalts between Scoresby Sund and Kangerlussuaq (Larsen et al., 1989; Neuhoff et al., 1997). Mineralogical zones define the postextrusive isograds in the area, are essentially uniform in thickness throughout the province and transgress lava stratigraphy. Neuhoff et al. (1997) studied the pronounced discordance between dips of zeolite isograds (c. 2.5°SE) and lava flows (4°–11°SE) in the coastal flexure zone and suggested that the tilting of the lava pile was pre-zeolite formation and syn-volcanic due to the progressive focussing of volcanism along the coast. Neuhoff et al. (1997) explained the pronounced dip of the isograds with post-metamorphic deformation associated with dikes

J.M. Bonow et al. / Global and Planetary Change 116 (2014) 10–29

and faults near the continental margin. According to Neuhoff et al. (1997), zeolitisation occurred rapidly during and just after the bulk of volcanism because basalt porosity was rapidly reduced after zeolite formation and thus limiting groundwater flow and further development of the zeolites. Isograds represent levels of equal thermal alteration and, according to Larsen et al. (1989), the absence of the shallow and less altered zeolite zones may be explained by the removal of these zones by erosion. The amount of section removed since the formation of the zeolite zones was shown to vary between 250 and 1850 m at six locations studied by Larsen et al. (1989) and by Neuhoff et al. (1997) (Table 1). We calculated the amount of erosion by subtracting the depths to the isograds below the present surface from the reference depths below the paleo-surface as established by Larsen et al. (1989) based on a geothermal gradient of 40 °C/km. The minimum amount of removed section would increase to 800 m if the paleo-geothermal gradient was only 30 °C/km (Larsen et al., 1989). The results show that minimal erosion (400 m) has taken place across the Geikie Plateau and Gronau Nunatakker. Larsen et al. (1989) noted that the lost lavas across the Geikie Plateau are mainly of the Skrænterne Formation and that this formation probably extended further in the direction towards Milne Land. The authors also entertain the possibility that there was an overlying sequence of alkaline lavas equivalent to what is now known as the Prinsen af Wales Bjerge Formation (Hansen et al., 2002) and that possible feeder dykes for such lavas are present in the area. However, the estimated erosion increases to 700 and 900 m towards the north across Gåseland and Milne Land in

19

agreement with the gradual truncation of the basalt units in that direction (Fig. 9; see Larsen et al., 1989). Consistent with this, Larsen et al. (1989) concluded that the zeolitisation had a regional character and that it took place after the tilting of the lava sequence and before erosion had planed the sequence down to its present level. Towards the southern limits of the main area of flood basalts, the amount of erosion based on the zeolite isograds becomes significant at Lindbergh Fjelde and Nansen Fjord and reaches 1250 and 1850 m, respectively. The amount of section removed increases by 900 m over the 35 km between Gronau Nunatakker and Lindbergh Fjelde. The increase agrees well with the dip of the basalts of 2° towards Gronau Nunatakker (Pedersen et al., 1997). According to Neuhoff et al. (1997) this cover consisted of parts of the Geikie Plateau Formation, the Rømer Fjord and Skrænterne Formations. 4.9. Post-volcanic structures Substantial, post-volcanic movements caused considerable changes in the geometry of the Main Basalts (e.g. Brooks and Nielsen, 1982; Brooks, 1985; Pedersen et al., 1997). Pedersen et al. (1997) identified these movements and mapped several tectonic blocks with dips varying from 1 to 2° N-NE in inland areas to 12°SE at the coast. Pedersen et al. (1997) identified a nearly horizontal inland plateau between Lindbergh Fjelde and Gronau Nunatakker (Fig. 6), with flow dips less than 2°. Remnants of mildly alkaline basalts on several high nunataks suggest that the area may have been covered by such lavas which have been removed by erosion.

Table 1 Zeolite isograds in southern East Greenland based on the work of Larsen et al. (1989) and Neuhoff et al. (1997). Estimates of amounts of sections removed and rock uplift (upper limit) after zeolitisation.

Gronau Nunatakker Summit 2700 m asl

Reference (A) Isograd depth b p–surf. (m) Zeolite free Chabazite + thomsonite Analcime Mesolite – scoleicite Heulandite + stilbite Laumonite Mean value

0 400 1400 1600 2300 2500

(B) Isograd depth b surf. (m)

(C)

(D)

(E)

Isograd elev. (m asl)

Rem. sect. (m)

Rock uplift (m)

100 1100

2600 1600

300 300

3000 3000

300

Zeolite free Chabazite + thomsonite Analcime Mesolite – scoleicite Heulandite + stilbite Laumonite Mean value

0 400 1400 1600 2300 2500

Nansen Fjord Summit 1250 m asl

(B) Isograd depth b surf. (m)

(C)

(D)

(E)

Isograd elev. (m asl)

Rem. sect. (m)

Rock uplift (m)

250 450 1000

2750 2550 2000

1150 1150 1300

4150 4150 4300

1200

4200

3000

Geikie Plateau Summit 1900 m asl

Reference (A) Isograd depth b p–surf. (m)

Lindbergh Fjelde Summit 3000 m asl (B) Isograd depth b surf. (m)

(C)

(D)

(E)

Isograd elev. (m asl)

Rem. sect. (m)

Rock uplift (m)

400 700

850 550

1900 1800 1850

3150 3050 3100

Milne Land Summit 1800 asl

Gåseland Summit 1800 m asl

(B) Isograd depth b surf. (m)

(C)

(D)

(E)

Isograd elev. (m asl)

Rem. sect. (m)

Rock uplift (m)

0 1000 1150

1900 900 750

400 400 450

2300 2300 2350

2200

–100

300 388

2400 2338

(B) Isograd depth b surf. (m)

(C)

(D)

(E)

Isograd elev. (m asl)

Rem. sect. (m)

700

1100

(C)

(D)

(E)

Rock uplift (m)

(B) Isograd depth b surf. (m)

Isograd elev. (m asl)

Rem. sect. (m)

Rock uplift (m)

700

2500

500

1300

900

2700

700

2500

900

2700

(A) Isograd reference depth below paleo-surface assuming a paleogeothermal gradient of 40 °C/km (Larsen et al., 1989). (B) Isograd depth below present surface. Data from Neuhoff et al. (1997) (Gronau Nunatakker, Lindbergh Fjelde, Nansen Fjord) or calculated as summit elevation − C. (C) Isograd elevation. Data from Larsen et al. (1989) (Geikie Plateau, Gåseland, Milne Land) or calculated as summit elevation − B. (D) Removed section = A − B. (E) Rock uplift = A + C. Upper limit assuming paleo-surface near sea level. b: below, elev: elevation, p-surf: paleo-surface, Rem. sect.: removed section.

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J.M. Bonow et al. / Global and Planetary Change 116 (2014) 10–29

Fig. 11. Erosion surfaces, escarpments and fault blocks in East Greenland. An etch surface (ES, see also explanation in text) is identified in the northern part of the study area (see Fig. 3A). The Upper Planation Surface (UPS) cuts the basement and the Paleogene Main Basalts, as well as across blocks tilted by post-basalt movements (e.g. Lindbergh Block, cf. Pedersen et al., 1997). This shows that the UPS was formed by erosion in post-basalt times. We also identify distinct blocks (e.g. the Gåseland Block) where the UPS is vertically offset relative to the adjacent areas. We interpret this as an indicator of tectonic movements after the formation of the UPS. The LPS extends over wide areas at the head of Kangerlussuaq, but it occurs only as valley benches on southern Milne Land and in the areas west of Milne Land. Geology in background, legend in Fig. 2. See online map for further details.

Brooks (1979, 1985) used geological and geomorphological evidence to introduce the concept of the “Kangerlussuaq dome” (Fig. 2). He argued that about 3 km of basalt had accumulated above the present-day summit of Domkirkefjeldet during the initial event of lava eruption and net subsidence and that the dome surface formed subsequently by uplift culminating around Domkirkefjeldet where the top of the basalts reached about 6 km a.s.l. Brooks used young apatite fission-track ages published by Gleadow and Brooks (1979) to show that the dome was already deeply eroded by 35 Ma and concluded that the dome most likely formed around the time of emplacement of the Kangerlussuaq Intrusion around 50 Ma. Brooks (1985, 2011) and

Pedersen et al. (1997) suggested that a large block occupying more than 1500 km2 in Watkins Bjerge, represents the flank of “the Kangerlussuaq dome”. Here the basalts reach up to 3.7 km a.s.l. and the basalts within the block tilt as much as 4°NE away from centre of the presumed “Kangerlussuaq dome”. The Ocean Drilling Program (ODP) Leg 152 off SE Greenland documented subaerially erupted lavas 3.1 km below seabed (~63°N; Larsen et al., 1993). Onshore southern East Greenland, the flood basalts erupted onto a largely horizontal lava plain near sea level (see above), but today these lavas reach 3.7 km a.s.l. at Gunbjørn Fjeld, the highest mountain in Greenland. Consequently, post-breakup, differential vertical

Fig. 12. Photos from the study area with index map of photo locations. A. Milne Land looking south-east across Scoresby Sund. The etch surface (ES) is developed in basement rocks and is characterised by distinct hills, up to 100 m high. In the foreground, the ES emerges from below a Jurassic cover, while on the other side of Scoresby Sund it is overlain by Paleogene basalt. The age of the ES is thus different on either side of Scoresby Sund, but shape alone cannot be used to distinguish the two surfaces from each other in areas where the cover has been stripped. The upper planation surface (UPS) cuts across the Paleogene basalt. The UPS extends south of Scoresby Sund where it is known as the Geikie Plateau. B. Central Milne Land. The UPS (blue dashed line) cuts across Paleogene basalt and forms a plateau at c. 1800 m a.s.l. The basement–basalt contact (red dashed line) is highly irregular and the basement is deeply weathered. This etch surface is re-exposed in the inner parts of Scoresby Sund. C. Milne Land looking north across Øfjord (note the icebergs). The UPS cuts across basement and forms a plateau at 1900 m a.s.l. D. Inner part of Kangerlussuaq looking north-east to the Lower Planation Surface (LPS, here known as the Nordfjord Plateau) which cuts across basement rocks and, further inland, across the Eocene basalts of the Prinsen af Wales Bjerge Formation (Hansen et al., 2002). A several hundred metre high escarpment, the Lindbergh Escarpment is seen in the far distance. It separates the LPS from the rather dissected UPS (developed across basalts) on the Lindbergh Block above the escarpment. The elevation of the LPS is about 2000 m a.s.l., while the UPS reaches about 3000 m a.s.l. E. Inner part of Kangerlussuaq looking south to alpine relief around Domkirkefjeldet. Alpine topography is the dominant landscape type along Blosseville Kyst and no remnants of planation surfaces are preserved in these landscapes. F. East side of Christian IV Gletscher looking north to the ~1.5 km high erosional Watkins Escarpment (formed in basalt) along the edge of Watkins Bjerge. The escarpment trends generally W–E but is winding showing that it is an erosional feature and not primarily controlled by a fault. G. Watkins Bjerge looking north. The ice-filled valley is part of a major fluvial paleo-drainage system that heads generally northwards, away from the present coast. The Wager Plateau and the Wager Escarpment are seen in the distance. The UPS is highly dissected in the foreground along the escarpment, but forms the extensive Wager Plateau in the distance.

J.M. Bonow et al. / Global and Planetary Change 116 (2014) 10–29

C

21

D UPS

B A

Lindbergh Escarpment LPS

DE

Nordfjord Plateau

G F

E

A Geikie Plateau (UPS)

Scoresby Sund

ES

Bas

eme nt

Jurassic

F

Basalt

Watkins Bjerge

Watkins Escarpment

B

UPS

LPS? Basalt

Basement ES

G Wager Plateau (UPS)

C

UPS

Wager Escarpment

22

J.M. Bonow et al. / Global and Planetary Change 116 (2014) 10–29

A 4 3 2 1 0

Gunbjørn Fjeld

(km) SSW

Watkins Escarpment Alpine relief Alpine relief

?

.

NNE Wager Plateau

UPS

Gåseland Block

UPS

IED

UTM Y

7575000

7600000

7625000

B 4 3 2 1 0

Watkins Fault Watkins Block Wager Escarpment UPS UPS

7675000

Lindbergh Escarpment

(km) SSW Skærgaard f Alpine relie Intrusion UTM Y

7650000

7575000

7600000

UPS

7700000

7725000

7775000

7800000

7825000

7850000

7875000

7900000

Maximum topography along swath

NNE

LPS

Topography 7625000

7650000

7675000

c

7700000

Minimum topography along swath

C (km) 4 WNW 3 2 1 0 835000

7750000

Fig. 9

ESE Milne Land

UPS

UPS

Jurassic

ES 860000

885000

910000

935000

960000

b

a

985000 UTM X

Fig. 13. Topographical profiles combined with maximum and minimum topography within swaths covering 25 km on each side of the topographical line. A. SSW–NNE profile between Nansen Fjord and Gåseland. The geology along the profile is dominantly basalt, except for north of point 7825000 where the bedrock is basement. The profile illustrates the presence of two tectonic blocks, the Watkins Block and the Gåseland Block which have moved vertically after formation of the UPS. The Watkins Escarpment is erosional and is a distinct landscape feature. B. SSW–NNE profile along the eastern side of Kangerlussuaq. The geology in the south is dominated by basement rocks and the Skaergaard Intrusion but, north of point 7640000 the profile intersects basalt. The profile illustrates the relationship between the LPS, the UPS and the erosional Lindbergh Escarpment that separates the two surfaces. No surfaces or surface remnants have been identified in the coastal areas where glacial erosion has resulted in alpine relief. C. WNW–ESE profile across Milne Land. The basement rocks on eastern Milne Land are capped by basalt in the summits and by Jurassic sediments on the eastern slope towards Scoresby Sund. The Jurassic cover, directly onto basement, constrains the minimum age for this deeply weathered ES in this area. The profile also illustrates the geometrical relationship between the near-horizontal UPS that cuts off the inclined ES. The UPS is therefore younger than the ES. The UPS is slightly offset across Rødefjord, which is probably related to movements of the Gåseland Block to the south. Upper Planation Surface (UPS): stippled blue line, Lower Planation Surface (LPS): stippled orange line, etch surface (ES): stippled red line. Numbers on the length axis refer to (m) in UTM zone 24. IED: area with inaccurate elevation data. Location of profiles indicated in the index map.

movements with a magnitude of more than 6 km have shaped the present-day geometry of the large igneous province along the East Greenland margin (Fig. 10). 5. Stratigraphic landscape analysis 5.1. Mapping of peneplains We identified extensive etch surfaces (ES) formed by the deep weathering of basement rocks and two other surfaces of regional extent and low relative relief that we identified as planation surfaces: the Upper Planation Surface (UPS) and the Lower Planation Surface (LPS) (Fig. 11 and map at the scale of 1:500,000 in the supplement online). In the following we show that the UPS and the LPS cut across rocks of different ages and resistances and that they therefore are erosional features. We also identify and name significant erosional escarpments (Lindbergh, Wager and Watkins Escarpments) and our mapping reveals that the surfaces, especially the UPS, are tilted along or offset across lineaments (Fig. 11). This observation is important because it shows that faults have been active after the formation of the UPS. We are thus able to identify faults (one is named here, Watkins Fault) and four tectonic blocks that are limited by such faults (Gåseland, Kap Dalton, Lindbergh and Watkins Blocks; Fig. 11). Finally, we identify pre-glacial drainage patterns and evidence for significant glacial erosion in some areas and formation of alpine relief (Figs. 12 and 13). 5.1.1. Etch surface (ES) The deeply weathered surfaces that formed prior to the eruption of the Paleogene basalts are important in the context of this study because they can be used to establish an erosional and depositional, pre-basalt history of the margin. However, in detail they can occasionally be identified as a pre-Middle Jurassic surface (e.g. on parts of Milne Land) or as a pre-mid-Paleocene surface (e.g. on Milne Land and Gåseland). However, a detailed analysis of these surfaces is beyond the scope of this paper and, as they are similar in character, we treat them as one etch surface (ES).

The ES that extends across the crystalline basement in the northern part of the study area is characterised by distinct hills that are a few hundred metres high and limited by narrow valleys along old fracture systems (cf. Birkelund and Perch-Nielsen, 1976; Larsen et al., 1989). Such “hilly relief surfaces” are formed in a warm, humid climate by deep weathering of the landscape, followed by stripping of the saprolite cover (e.g. Lidmar-Bergström et al., 1997; Bonow, 2005). The ES is recognized over wide areas of Milne Land and Gåseland (Figs. 3, 11 and 12) and also along the southern coast of Scoresby Sund, where basement is exposed below the more widespread Main Basalts. The hilly relief characteristics of the basement can be observed both in areas without cover rocks and also in areas where the basement/cover rock contact can be followed as an unconformity in exposed hillsides in the landscape (Fig. 3A). The ES occurs at different elevations in the landscape, from sea level up to 1800 m a.s.l. (Larsen et al., 1989), where the ES is cut off by the UPS (Figs. 12 and 13). In areas near cover rocks, the crystalline basement has a glacial overprint, for example striations and chattermarks. This shows that glacial erosion has played a significant role in the stripping of the Mesozoic–Paleogene cover.

5.1.2. Upper Planation Surface (UPS) We have mapped the UPS across an area of about 75,000 km2 between Kangerlussuaq and Scoresby Sund where a significant part of the plateau is known as the Geikie Plateau. A previously unnamed plateau area we name the Wager Plateau (Figs. 11 and 13). The elevation of the UPS is between 2 and 3 km a.s.l. over most of the study area (Fig. 14). In general, the UPS is best preserved inland while it is severely dissected or totally obliterated in the coastal areas. We infer that the UPS exists beneath the many local icecaps which cover a significant part of the study area. To reach that conclusion, we have complemented the identification of the UPS on digital elevation data by inspection of aerial photographs to identify the flat rock surfaces that emerge along the rim of the ice-caps (Fig. 3). In other cases, nunataks with minor remnants of flat surfaces occur and such areas are marked as “UPS inferred” on the map in Fig. 11.

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Fig. 14. Elevations and the general trend of elevation of the UPS and LPS along Blosseville Kyst. The UPS is coherent across wide areas and is generally tilted towards the north with the highest elevations around Christian IV Gletscher and the Watkins Block. The LPS extends over wide areas at the head of Kangerlussuaq but only as elevated valley benches around Milne Land. The red question marks indicate areas where the coherent elevation of the summits indicate that they are near the LPS, however, these areas are dominated by alpine relief without surface remnants. UPS: upper planation surface. LPS: lower planation surface. Location of profiles in Fig. 13 is shown.

The UPS can be followed from profile to profile across different rock types (e.g. basement and basalt on Milne Land) (Fig. 7), so the UPS is not controlled by lithology. The surface cuts across rocks of different ages, in particular across the Caledonian basement and the Main Basalts (Figs. 12 and 13). In some areas, there is hardly any angular unconformity between the UPS and underlying basalt flows (e.g. Gronau Nunatakker; Fig. 11) whereas there is a pronounced unconformity along a profile from the Geikie Plateau to Milne Land (Fig. 9; Larsen et al., 1989; Pedersen et al., 1997). Around Gåseland and Milne Land, the UPS also truncates the hilly relief (Fig. 12A).

5.1.3. Lower planation surface (LPS) The LPS is well-defined across the basement and basalt along the inner part of Kangerlussuaq at about 2 km a.s.l. (Figs. 11 and 12D). The surface can be mapped with high confidence only in a rather small area, c. 50 km by 50 km and includes the Nordfjord Plateau. Towards NE, the Lindbergh Escarpment is a clear feature that separates the LPS from the UPS (Fig. 12D), although there are 10–30 km between the area where the LPS is exposed and the escarpment that borders the UPS. There are only few nunataks in the intervening area which is almost entirely covered by ice. We mapped some areas as “inferred LPS” where there is a lack of fairly large flat areas, but where correlations between profiles allow connection between small flat surface occurrences. The low elevation of the LPS compared to the UPS and the erosional escarpments between the UPS and the LPS, show that the LPS formed by erosion below the UPS. A consistent level of summits in the profiles indicates that the LPS may have extended further east along the escarpment bordering

Watkins Block and the Wager Plateau, for example east of Christian IV Gletscher and along the present coast, from Kangerlussuaq to Scoresby Sund (Figs. 8, 11 and 14). We did, however, not identify any flat surfaces in these areas, whereas the LPS is readily identified in the area around Milne Land (Figs. 3 and 11), where it forms valley benches at approximately 1 km a.s.l. 5.1.4. Erosional escarpments 5.1.4.1. Wager Escarpment. The Wager Plateau (part of the UPS) is limited on its southern flank by a c. 1 km high escarpment, the Wager Escarpment (Figs. 11 and 12G). We interpret this escarpment as primarily erosional as it is winding, but its origin is probably scarp retreat from the east–west trending fault, the Watkins Fault, that offsets the Wager Plateau relative to the Watkins Block. 5.1.4.2. Watkins Escarpment. The southern flank of Watkins Bjerge forms a major and exceptionally high (1.5 km) escarpment, the Watkins Escarpment (Figs. 11 and 12F). The escarpment separates the summit level (near the UPS) from a lower level in the terrain which we infer to be near the LPS (see above). We interpret the escarpment to be an erosional feature as there is no change in lithology and because it is winding and thus that it is not fault controlled. It was most likely formed by scarp retreat after movements along the Christian IV Fault (cf. Larsen and Whitham, 2005). 5.1.4.3. Lindbergh Escarpment. The escarpment NE of Kangerlussuaq, the Lindbergh Escarpment, is less pronounced, but it rises a few hundred

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metres above the ice-covered plateau of the LPS and is clearly visible in the landscape (Figs. 11 and 12D). The escarpment is in many parts abrupt, but the high degree of dissection and its highly irregular appearance makes it sometimes difficult to establish its precise position. It is most likely erosional, similar to the Watkins Escarpment and to the Wager Escarpment, as the same type of basalt is present on both sides of the escarpment and it thus does not reflect differences in bedrock resistance. 5.1.5. Tectonic and erosional features 5.1.5.1. Gåseland Block. Gåseland Block is located on the innermost part of the Gåseland peninsula (Figs. 3A, 11 and 13). The block is downfaulted along the inner part of Rødefjord compared to the remnants of UPS identified on the NW side of the fjord. It is possible that Milne Land to the north constitutes another fault block. 5.1.5.2. Kap Dalton Block. Kap Dalton Block is situated along the northern part of Blosseville Kyst, between the coastline and Steno Bræ (glacier) and is 25–30 km wide. The drainage pattern of this glacier is peculiar as it runs parallel to the coast for more than 100 km. Kap Dalton Block is slightly tilted towards the coast and the summit may be part of the UPS (Figs. 3B, 10 and 11). 5.1.5.3. Lindbergh Block. Lindbergh Block lies west of Watkins Block and it is bordered by the Christian IV Gletscher in the east (Figs. 11, 12D and 13B). It is, like Watkins Block, limited by a fault to the north (maybe a continuation of the Watkins Fault), but the escarpment there is not as pronounced as Watkins Fault limiting the northern side of Watkins block. The southern flank of the Lindbergh Block is limited by the erosional Lindbergh Escarpment (Fig. 12D). It is possible that the Lindbergh and the Watkins Blocks constitute one single block that has been dissected by the Christian IV Gletscher. 5.1.5.4. Watkins Block. Watkins Block contains the high summit of Gunbjørn Fjeld and it is the most salient topography within the study area (Figs. 11, 12F and 13). The block consists of Watkins Bjerge and is well defined on its northern side along a west–east trending glacial valley, which we interpret to follow a fault, the Watkins Fault. However, the Watkins Fault seems to fade towards the west, as high topography is also present north of the west–east trending glacial valley. The steep, erosional Watkins Escarpment limits the block to the south. The limit of the block is poorly defined in the east while it is clearly controlled by Christian IV Gletscher with the inferred fault (Larsen and Whitham, 2005) in the west. We interpret the summits of the Watkins Block to be near the UPS because they can be correlated with the UPS on the Wager Plateau via the profiles. These observations imply that Watkins Block moved significantly in post-UPS time. We have, furthermore, identified and mapped a paleo-drainage system at high elevations in the area around Watkins and Lindbergh Blocks (Figs. 11 and 12G). These patterns of fluvial valley systems were noted by Brooks (1985), but not mapped in detail. We note that the paleodrainage points in a generally northwards direction, away from the present coast (Fig. 11) and that it has an ordinary, fluvial dendritic pattern. At present these valleys are filled with glaciers flowing northwards and the original valleys have been deepened and widened. The drainage direction eventually changes towards the south as the valleys connect to the southward-flowing Christian IV Gletscher that drains into the Atlantic at the southern edge of Blosseville Kyst (Fig. 11). This flow pattern shows that the original drainage pattern has been altered and that the water divide has moved from a position closer to the present coast to a more inland position, further north-west. The mountains in the coastal regions are dominated by alpine relief characterised by glacial erosional forms such as cirques, horns and arêtes and deep glacial valleys. This erosion has resulted in total destruction of any preglacial plateaux.

5.2. Denudation chronology and magnitude of uplift 5.2.1. Constraints on the timing of the formation of the UPS and the LPS The UPS and LPS are regionally extensive, low-relief surfaces that cut across rocks of different ages and resistance, in particular across the Main Basalts. We thus interpret them to be erosion surfaces (peneplains; cf. Green et al., 2013 and references therein) and infer that they were graded to the general base level during the time of their formation. This must have happened after the extrusions of the basalts, after the onset of sea-floor spreading in the NE Atlantic and after an event of regional tilting of the basalt flows (cf. Fig. 9; Pedersen et al., 1997). Our preferred interpretation is that the base level to which these surfaces were graded was the level of the adjacent Atlantic Ocean. The stratigraphy off South-East Greenland (c. 63°N) provides further insight into the denudation chronology of the margin (Ocean Drilling Project, leg 152, site 918; Larsen et al., 1994b). Larsen et al. (1994b) thus found that the lower Eocene sediments at site 918 indicated low sedimentation rates with limited terrigenous influx before a middle Eocene–upper Oligocene hiatus, whereas the sudden and strong influx of coarse clastic turbidites at this site during the late Oligocene may have been triggered by an uplift of the inner margin. These observations indicate that a mid-Cenozoic uplift event caused regional tilting of the basalts, deep erosion of the margin and ultimately, formation of the UPS and LPS (see also Larsen and Saunders, 1998). The uplift of the margin that led to the formation of the UPS thus occurred between about 45 Ma and about 30 Ma. The regional extent of the UPS between Kangerlussuaq and Scoresby Sund (Figs. 3, 11 and 12) suggests that the surface formed during a long time span dominated by stable baselevel conditions. It is possible to further constrain the timing for the formation of the UPS and LPS due to the presence of the middle Miocene lava flows of the Vindtop Formation (14–13 Ma) that overlie the Main Basalts in a small area along the southern rim of the Wager Plateau (Figs. 2 and 11; Storey et al., 2004; S. Watt, pers. com. 2012). Two observations are essential: first, that the Vindtop Formation (2.7–2.9 km a.s.l.) is located above the UPS (mapped at 2.6 km a.s.l. near the outcrop of the Vindtop Fm, Fig. 14). Second that both the UPS and the LPS are regional in extent and thus needed significant time to develop. In West Greenland the regional UPS peneplain needed 20 Ma to develop (Japsen et al., 2006) so we find it unlikely that both UPS and LPS could have developed after the middle Miocene (in which case the Vindtop Formation would be below the UPS). We therefor conclude that the Vindtop Formation erupted onto the UPS, and thus that UPS formed after 45 Ma and before 14 Ma. Furthermore, our preferred interpretation is that the uplift event that initiated the formation of the LPS most likely happened after the eruption of the Vindtop Formation. There are, however, no geological constraints confirming this age, but the age of the LPS is constrained to be younger than the UPS, and the regional extend of the UPS implies that significant time must have been available for its formation and thus that the formation of the more restricted LPS must have started relatively late. Thus when estimating the time interval for the formation of the LPS, some key aspects must be accounted for; namely: (1) the time to produce the LPS; (2) a phase of regional uplift of the LPS; and (3) time to destruct the uplifted LPS, creating alpine relief and incised valleys. A tighter time constrain for the LPS is offered by integrating with AFTA data (Japsen et al., 2014–in thi issue). 5.2.2. Estimation of the amount of uplift Our interpretation of the UPS and the LPS as erosion surfaces graded to former sea levels implies that their present elevation can be used to estimate the amount of net uplift since their formation. The amount of uplift is equal to the vertical distance between the present elevation of the UPS or LPS relative to the present base level; i.e. the sea. The presence of two elevated erosion surfaces across southern East Greenland, moreover, implies that phases of uplift affected the margin in postUPS time; i.e. since the middle Miocene. Each uplift phase inevitably

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led to valley incision by rivers, initiating new phases of planation towards sea level; first the LPS and then the present-day valleys. The vertical distance between the UPS and the LPS is thus an estimate of the amount of uplift in the first phase of post-UPS uplift whereas the elevation of the LPS provides an estimate of the uplift in the second phase. Our results show that the magnitude of the first phase of uplift was about 1 km over most of the study area, but that the magnitude was about 1.5 km along the southern side of Watkins Block (Fig. 14). The second phase had a magnitude of 2 km in the Kangerlussuaq area and 1 km around Milne Land. 5.2.3. Relative denudation chronology As demonstrated in Section 4, the land surface of southern East Greenland remained near sea level during the eruption of the Main Basalts and the character of the UPS and the LPS is best explained by these surfaces being graded to base level at sea level during the time of their formation. Miller et al. (2005) showed that eustatic sea-level fall has been less than 100 m since 50 Ma, so the inferred changes of surface elevation since breakup are much larger than can be explained by a lowered sea level change only. Medvedev et al. (2013) analysed the influence of the ice-sheet loading in the central part of Greenland and the carving of the fjord systems on the evolution of the topography by numerical modelling processes backward in time. They found that peripheral bulging caused by the load of the Greenland ice sheet is insignificant whereas glacial carving can cause uplift of up to 1.2 km. Medvedev thus concluded that much of Greenland's topography is not caused by icerelated processes and that the origin of the pre-glacial mountain chains was enigmatic. A tectonic trigger is therefore needed to explain the present relief. We suggest that the formation of the planation surfaces and their present elevation can best be explained with the following relative chronology of events (Figs. 5 and 6): (1) Denudation and weathering of the basement in pre-Middle Jurassic and pre-mid-Paleocene times led to the formation of deeply-weathered basement surfaces characterised by hilly relief (ES). Rifting, subsidence and burial in the Cretaceous to Paleocene generated the Kangerlussuaq Basin. (2) Short-lived uplift in the mid-Paleocene due to the impact of the Iceland plume caused rotation and erosional truncation of the sediments below the earliest flows of the Lower Basalts that were partly extruded in a marine environment. (3) Breakup of the NE Atlantic at the Paleocene–Eocene transition at c. 56 Ma was accompanied by the extrusion of the Main Basalts during rapid subsidence (Fig. 6A). (4) Syn-volcanic tilting of the lava sequence took place due to progressive focusing of volcanism at and beyond the coast. Subsidence and burial during mid-Eocene eruption of the late volcanics (the Igtertivâ Formation with marine incursions) and deposition of sediments (fluvial to shallow-marine deposits of the Kap Dalton Group) continued. (5) Mid-Cenozoic uplift (sometime between 45 and 30 Ma), regional tilting and erosion led to the formation of the UPS near sea level before the middle Miocene (Fig. 6B). (6) The mid-Miocene Vindtop Formation extruded onto the UPS (Fig. 6B). (7) Post-mid-Miocene uplift of the UPS by c. 1 km across most of the study area resulted in valley incision below the UPS, dissection and tilting of the UPS and subsequent formation of the LPS near sea level. Watkins Block and Lindbergh Block were uplifted by up to 1.5 km along their southern edges in this phase and the northward tilting of the blocks induced drainage in that direction (Fig. 6C, D). Uplift of the UPS and LPS by c. 2 km in the Kangerlussuaq area and c. 1 km in the Milne Land area followed by valley incision and glacial erosion led to the formation of the present-day landscape (Fig. 6E).

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6. Discussion 6.1. Formation and preservation of planation surfaces The UPS is well preserved in areas with resistant basement rocks (e.g. Milne Land) and also in areas covered with basalt at some distance from the coast (e.g. the Wager Plateau). This contrasts with West Greenland where the equivalent UPS is best preserved in basement rocks, while basalt areas are much more dissected (e.g. on Nuussuaq; Bonow et al., 2006a). The UPS in East Greenland is at significantly greater elevations than in West Greenland. The higher elevations may also be the explanation for the better preservation of the UPS in East Greenland because it is well established that landscapes, sometime called “relict surfaces”, are preserved beneath non-erosive cold-based ice (Kleman, 1994; Kleman and Hättestrand, 1999). The LPS is best preserved in areas of basement rocks (e.g. the Nordfjord Plateau), while areas where we were able to infer only the presence of the LPS, coincide with areas covered by basalts or sedimentary rocks that are at lower elevation and thus more prone to experience wet based and erosive glacial conditions. These areas are also close to the present coast, which favours the development of cirque glaciers with subsequent formation of alpine relief. Similar landscape patterns are observed in West Greenland (Bonow et al., 2006a,b) and in western Scandinavia (Lidmar-Bergström et al., 2000, 2007; Hall et al., 2013). 6.2. Estimation of removed section and rock uplift from zeolite isograds We have concluded that the UPS was formed after a phase of uplift and erosion that followed the deposition of Paleogene volcanics and sediments. Zeolite isograds represent levels of equal thermal alteration and the absence of the shallow and less altered zeolite zones may be explained by the removal of these zones by erosion (Larsen et al., 1989). Estimates on the amount of erosion based on the absence of zeolite zones (Table 1) thus provides a lower limit on the amount of section removed during formation of the UPS. The UPS is well-defined across Geikie Plateau and over Gåseland and Milne Land and the estimated amounts of erosion based on zeolite zones (400 and 700–900 m, respectively) are thus minimum estimates of the section removed in these areas during the formation of the UPS. At Lindbergh Fjelde, the UPS coincides with the summits and this allows us to conclude that this surface formed after tilting of the basalts and removal of a cover of a minimum of 1250 m. Nansen Fjord is in the coastal zone where the UPS cannot be identified. We can use the zeolitisation (which happened pre-UPS) to estimate an upper limit for net rock uplift. The upper limit case corresponds to a situation where the paleo-surface during zeolite formation was close to sea level and thus to net subsidence during the eruption of the Main Basalts (Table 1). In this model, rock uplift equals summit elevation plus the amount of removed section; e.g. Gronau Nunatakker 3000 = 2700 + 300 m. The upper limit of rock uplift varies from 2.3 km at Geikie Plateau to 4.2 km at Lindbergh Fjelde. 6.3. The “Kangerlussuaq dome” and other structures Dome-shaped topographic features are common in recently uplifted areas such as southern Norway (the Southern Scandes) and southern Sweden (the South Swedish Dome, e.g. Lidmar-Bergström et al., 2000, 2013) and Nuussuaq, West Greenland (Bonow et al., 2006a). In our study area the “Kangerlussuaq dome” (Fig. 2) has been discussed in several papers (Wager, 1947; Brooks, 1979, 1985, 2011; Pedersen et al., 1997). However, the “Kangerlussuaq dome” as identified by Brooks (1979, 1985) is not a topographical dome in the present landscape, but rather a geological structure, formed by the complex interaction of uplift and erosion since the emplacement of the Kangerlussuaq Intrusion in early Eocene (Figs. 2 and 5). The present topography in our

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study area is instead dominated by near-horizontal erosion surfaces and by deeply incised valleys. Previous studies have suggested that Watkins Bjerge (and hence Watkins Block) is at the northern edge of the “Kangerlussuaq dome” (Brooks, 1985, 2011; Pedersen et al., 1997). We want to point out that the high elevation of these mountains is unrelated to the early Eocene land surface and that the present topography is the combined result of: (1) a first uplift event leading to planation and formation of the UPS by truncation of all Paleogene structures; (2) a second event leading to uplift of the UPS, block tilting and erosion to form the LPS south of the Watkins Escarpment and finally (3) a third uplift event leading to the elevation of the present landscape. The 4° tilting of the lava flows towards the north in the Watkins Block (Pedersen et al., 1997) is consistent with our interpretation that the summits here once were a continuation of the Wager Plateau. The UPS dips slightly towards the north in the area of Lindbergh Fjelde–Gronau Nunatakker, but the dip is steeper in the Lindbergh Fjelde area (the Lindbergh Block; Fig. 11). This interpretation is consistent with the results of Pedersen et al. (1997), who showed that the basalts are virtually horizontal around Gronau Nunatakker whereas the basalts are tilted 2° in Lindbergh Fjelde. Brooks (e.g. 2011) considered the Nordfjord Plateau at the head of Kangerlussuaq to be a pre-basalt surface, but our mapping shows that this plateau coincides with the LPS which also truncates the basalts at Prinsen af Wales Bjerge. 6.4. Drainage patterns It is well known from analyses of the pattern of the large valley systems in southern Norway that the water divide has migrated towards the east, away from the present coast (e.g. Ahlmann, 1919). At high elevations (c. 1500 m a.s.l.), broad and shallow valleys slope away from the Norwegian coast and cross the present water divide. However, migrating knick points from west-flowing rivers that end in the North Atlantic have successively captured the east flowing rivers (which have lower gradients). The river courses therefore form agnor valleys (fish-hook valleys) (e.g. Reusch, 1901; Ahlmann, 1919; Bonow et al., 2003 and references therein). The conjugate margin of East Greenland, especially in the Kangerlussuaq–Scoresby Sund area shows a similar valley pattern. Well-developed agnor valleys characterise the drainage pattern on the Lindbergh and Watkins Blocks (Fig. 12G). These old patterns of fluvial valley systems were noted by Brooks (1985) who thought that they had developed as superimposed drainage in response to the uplift of the “Kangerlussuaq dome”. However, due to the poor quality of the maps available to him, Brooks was unable to carry out a detailed analysis. Our favoured explanation for these northward-flowing systems and the formation of the agnor valleys is that the water divide was previously located close to the present coast and that it has migrated northwards due to late uplift. We thus speculate that the initial water divide that formed after the first uplift of the UPS is now lost due to post-UPS erosion and thus that the maximum post-UPS uplift was south of Watkins Block. 6.5. Timing and magnitude of tectonic events We used AFTA data to identify three Cenozoic events of uplift and exhumation in central West Greenland, starting between 36 and 30 Ma, between 11 and 10 Ma and between 7 and 2 Ma (Japsen et al., 2006). We were also able to correlate these events to the development and subsequent uplift of two regional planation surfaces (also named UPS and LPS). The evidence presented here does not allow us to establish the timing of uplift in detail, but the relative denudation chronology for southern East Greenland that we established from the stratigraphic landscape analysis overlaps with that for West Greenland (see Section 2).

We find that the UPS in southern East Greenland developed between 45 and 14 Ma whereas the timing for the formation of this surface in West Greenland was between ~35 and ~10 Ma. Thus the formation of the UPS in both West and East Greenland could be synchronous. Significant time with a relatively stable base level was thus available to form the UPS in both West and East Greenland. The landscape in West Greenland was at a low elevation until ~10 Ma when an uplift of 1 km of the UPS led to the incision of the LPS and formation of the major valleys. The uplift of the UPS in East Greenland and subsequent formation of the LPS happened after 14 Ma—again comparable to West Greenland. The LPS in southern East Greenland, extends across larger areas of the LPS in central West Greenland. This indicates either that a longer time was available for the formation of this surface in the east or that the formation of the surface was faster due to the presence of less resistant rocks (see above). The final phase of uplift in West Greenland was in the latest Neogene and a similar timing for uplift in East Greenland seems likely; at least we can say that it happened several millions of years after the onset of formation of the LPS that began after 14 Ma. Glacial erosion, with enhanced erosion in valleys (e.g. Sugden, 1974) has certainly contributed to isostatic compensation and further uplift of the surfaces, but cannot explain the full amount of uplift (Medvedev et al., 2008). 6.6. Comparison with the Norwegian margin A wide range of geological observations show that the margin of southern East Greenland was low-lying at the time of breakup and thus that the present-day relief is a consequence of events that happened long after breakup, a conclusion that has been reached by previous studies in East Greenland (e.g. Brooks, 1979, 1985, 2011; Bott, 1987; Larsen, 1990; Larsen and Marcussen, 1992; Clift et al., 1998; Larsen and Saunders, 1998). It is, however, only the present study that has clarified that the elevated plateaux in southern East Greenland and the deeply incised valleys below the plateaux are the result of three uplift phases that happened millions of years after breakup and that the presentday mountains were shaped after the mid-Miocene. This is interesting because elevated plateaux and incised valleys also characterise the large-scale landscape part of the conjugate margin in Norway (e.g. Lidmar-Bergström et al., 2000, 2013). The similarity of the two margins, both characterised by stepped surfaces, suggests that it is likely that the mountains of Norway may also have reached their present elevation in the late Cenozoic, long after the Atlantic breakup (cf. Bonow et al., 2007). Previous studies of the large-scale landscapes in Norway (e.g. Reusch, 1901; Ahlmann, 1919) have reached similar conclusions, i.e. that the plateau surfaces there (the paleic relief; Lidmar-Bergström et al., 2000) were formed by erosion to base level during the Cenozoic and that the surfaces reached their present elevation after uplift in the late Cenozoic (Lidmar-Bergström et al., 2013 and references therein). The absence of cover rocks younger than Devonian on the crystalline basement in these mountains (Sigmond et al., 1984), makes the timing of these events difficult to determine. Many geoscientists who are unfamiliar with the principles of large-scale geomorphology and stratigraphic landscape analysis have, therefore, remained sceptical (e.g. Nielsen et al., 2009a). Steer et al. (2012) estimated that 300–400 m was eroded off the high-elevation, low-relief surfaces in mid- and south Norway by glacial erosion during the Pliocene and Quaternary. On the contrary we find that glacial erosion in East Greenland has limited effect on the surfaces, but is focussed along the fjords. The surfaces are mainly “relict surfaces” and their preservation can most likely be explained by non-erosive, cold-based ice concluded from many studies of plateau surfaces in Scandinavia and North America (cf. Sugden, 1978; Rea et al., 1996; Kleman and Stroeven, 1997; Kleman and Hättestrand, 1999; Hall et al., 2013). The presence of the basaltic flows of the mid-Miocene Vindtop formation at high elevation, above the UPS, documents that this regional

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erosion surface took place before the onset of Neogene glacial activity in Greenland (Larsen et al., 1994a). The clear geological evidence from southern East Greenland combined with the analysis presented in this paper now documents that the mountains along this margin are young, yet pre-glacial in origin and we suggest that this is also the case for the similar landscapes of Norway. 6.7. Wider implications The results presented here clearly show that the East Greenland margin is not a remnant of a former orogen (e.g. Pedersen et al., 2012), nor a remnant of a rift shoulder or other structures related to rifting and break-up (e.g. Swift et al., 2008), but was shaped tens of millions of years after breakup and, as demonstrated by Medvedev et al. (2013), a topographic feature much of which already existed prior to the Pleistocene even though glacial carving caused additional uplift. Japsen et al. (2014–in this issue), in a paper accompanying the present found that the three phases of uplift in southern East Greenland are synchronous with phases in West Greenland, that they overlap in time with similar events in North America and Europe and that they also correlate with changes in plate motion. They also found that the much higher elevation of East Greenland compared to West Greenland suggests dynamic support in the east from the Iceland plume and that these observations thus indicate a connection between uplift along a passive continental margin, mantle convection and changes in plate motion. Japsen et al. (2012a) and Green et al. (2013) listed many elevated passive continental margins around the world that are characterised by elevated plateaux at 1 to 2 km or more a.s.l. cut by deeply incised valleys and are commonly separated from an adjacent coastal plain by one or more escarpments. Mesozoic–Cenozoic rift systems parallel to the coast are commonly present offshore with a transition from continental to oceanic crust further offshore. Japsen et al. (2012a) and Green et al. (2013) also noted that landscapes that characterise elevated passive continental margins are similar despite different geological settings and despite the time span since breakup (e.g. Japsen et al., 2012b). In particular, they noted that the landscape along the West Greenland margin, with its characteristic high-level plateaux, is not a remnant of the rifting process but is much younger (Japsen et al., 2005, 2006, 2009; Bonow et al., 2006a,b). Since the West Greenland margin shares all the characteristics of elevated passive continental margins as described above, the results from West Greenland led Japsen et al. (2012a) and Green et al. (2013) to suggest that elevated passive continental margins have formed as a result of an episodic development involving post-breakup subsidence and burial followed later by uplift and denudation. We can now add East Greenland to the list of elevated passive continental margins for which a post-rift history of episodic burial and exhumation is well documented. An immediate consequence of this insight is for geodynamic modellers to develop a theory that can explain elevated passive continental margins which is consistent with the observations highlighted here (cf. Cloetingh et al., 1990; Pedoja et al., 2011; Yamato et al., 2013). The results of the present study also underlines the importance of mapping and analysing erosion surfaces along elevated passive continental margins and of using them as tectonic markers (unconformities) for deciphering the development of such margins. Finally, the results underline the importance of episodic, vertical movement along elevated passive continental margins as highlighted by Green et al. (2013). 7. Conclusions We have applied stratigraphic landform analysis to identify three denudation surfaces that are important markers for the tectonic development of the East Greenland margin; etch surfaces (ES) and the Upper and Lower Planation Surfaces (UPS and LPS). We were able to use the geological conditions of the margin to conclude that its present

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elevation–and those of the UPS and LPS that define a landscape in two distinct steps–is the result of three regional phases of uplift and incision that happened long after breakup in the NE Atlantic. The UPS was formed by erosion to base level as a consequence of a first phase of uplift that happened after extrusions of the Paleogene flood basalts on a lava plain near sea level. The geological record offshore of SE Greenland constrain this phase to having happened around the Eocene–Oligocene transition. Our mapping shows that the volcanics of the middle Miocene Vindtop Formation accumulated on the UPS. The LPS formed by incision below the UPS after the second phase of regional uplift that most likely took place after the middle Miocene. This phase lifted the UPS by c. 1 km across most of the study area. The third phase of regional uplift must have taken place several million years after the second phase. This phase lifted the UPS and the LPS to their present elevations of 2–3 km and 1–2 km a.s.l., respectively These uplift events are of a magnitude that cannot be explained by isostatic compensation or eustatic sea level changes; they require a tectonic component. Given the similarity of the East Greenland landscapes with those on the conjugate margin in Scandinavia, we suggest that the mountains of Norway were also formed long after Atlantic breakup. Acknowledgements The research and publication was made possible by the IPY project “Mountain building and ice-sheet stability in Greenland” (Miss Green) which was funded by the Commission for Scientific Research in Greenland as part of the International Polar Year activities. We thank Kent Brooks, James A. Chalmers, Lotte M. Larsen, Asger K. Pedersen and Stuart Watt for invaluable input. We are very grateful to the comments from Guido Giordano and Laurent Husson on an early draft of the manuscript. We also want to thank Julien Babault and one anonymous person for positive and constructive review. The original data of ASTER GDEM is the property of METI and NASA and is publicly available at http://www.gdem.aster.ersdac.or.jp/index.jsp. The paper is published with permission of the Geological Survey of Denmark and Greenland. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.gloplacha.2014.01.010. References Ahlmann, H.W., 1919. Geomorphological studies in Norway. Geogr. Ann. 1, 1–148. Ahlmann, H.W., 1941. Studies in North-East Greenland 1939–1940. Geogr. Ann. 23, 145–209. Ahnert, F., 1998. Introduction to Geomorphology. Arnold, London. Amante, C., Eakins, B.W., 2009. ETOPO1 1 Arc-minute Global Relief Model: Procedures, Data Sources and Analysis. US Department of Commerce, National Oceanic and Atmospheric Administration, National Environmental Satellite, Data, and Information Service, National Geophysical Data Center, Marine Geology and Geophysics Division, Boulder (http://www.ngdc.noaa.gov/mgg/global/global.html). Babault, J., Van Den Driessche, J., Bonnet, S., Castelltort, S., Crave, A., 2005. Origin of the highly elevated Pyrenean peneplain. Tectonics 24, TC2010. http://dx.doi.org/ 10.1029/2004TC001697. Bengaard, H.J., Henriksen, N., 1984. Geology, Scoresby Sund. Map sheet no. 12, 1:500 000. Geological Survey of Greenland, Copenhagen. Birkelund, T., Perch-Nielsen, K., 1976. Late Palaeozoic–Mesozoic evolution of central East Greenland. In: Escher, A., Watt, W.S. (Eds.), Geology of Greenland. Geological Survey of Greenland, Copenhagen, pp. 305–339. Birkenmajer, K., 1972. Report on investigations of Tertiary sediments at Kap Brewster, Scoresby Sund, East Greenland. Rep. Geol. Surv. Greenl. 48, 85–91. Birkenmajer, K., Jednorowska, A., 1997. Early Oligocene foraminifera from Kap Brewster, East Greenland. Ann. Soc. Geol. Pol. 67, 155–173.

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