THE ROLE OF LITHOSPHERIC DELAMINATION AND ICE-DRIVEN ROCKFALL
EROSION IN THE EVOLUTION OF MOUNTAINOUS LANDSCAPES
by
TRISTRAM CHARLES HALES
A DISSERTATION
Presented to the Department of Geological Sciences and the Graduate School of the University of Oregon in partial fulfillment of the requirements for the degree of Doctor of Philosophy
December 2006 ii
“The Role of Lithospheric Delamination and Ice-Driven Rockfall Erosion in the
Evolution of Mountainous Landscapes,” a dissertation prepared by Tristram Charles
Hales in partial fulfillment of the requirements for the Doctor of Philosophy degree in the
Department of Geological Sciences. This dissertation has been approved and accepted by:
______Joshua J. Roering, Chair of the Examining Committee
______Date
Committee in Charge: Joshua J. Roering, Chair Eugene D. Humpherys Ray J. Weldon II W. Andrew Marcus
Accepted by:
______Dean of the Graduate School iii
© 2006 Tristram Charles Hales iv
An Abstract of the Dissertation of
Tristram Charles Hales for the degree of Doctor of Philosophy in the Department of Geological Sciences to be taken December 2006
Title: THE ROLE OF LITHOSPHERIC DELAMINATION AND ICE-DRIVEN
ROCKFALL EROSION IN THE EVOLUTION OF MOUNTAINOUS
LANDSCAPES
Approved: ______Joshua J. Roering
This dissertation discusses the evolution of mountains, particularly the interaction between uplift, which is controlled by horizontally-directed plate tectonic and vertically- directed isostatic forces, and erosion, which encompasses glacial, periglacial, fluvial and hillslope processes.
Eruption of the Columbia River Basalts (CRB) in northeastern Oregon is coincident with rapid uplift of the Wallowa Mountains. I mapped the modern distribution of CRB flows to quantify the amount of post-eruptive uplift in northeastern Oregon, which creates a broad “bull’s eye” pattern centered on a large granitic pluton. Rapid
Wallowa Mountain uplift appears to be related to delamination of dense lower crust beneath the Wallowa batholith and is the likely cause of flood basalt volcanism at ~17
Ma B.P. v
Rapid rockfall erosion rates in high mountains create extensive talus slopes. Parts of the Southern Alps, New Zealand, are dominated by kilometer-scale scree-mantled slopes which reduce hillslope gradients and decrease drainage densities. Field and aerial photograph-based analysis of scree slopes (i.e. rockfall deposits) reveal a peak in the areal extent and rate of scree slope formation in the eastern Southern Alps. This spatial distribution precludes earthquakes, post-glacial stress release, and rock type as primary rockfall triggering mechanisms. Instead, scree slopes occupy a narrow elevation range across the New Zealand Alps and the mean elevation of scree slopes is consistent with a theory for rockfall initiation by frost cracking and segregation ice growth, a process that breaks rocks through surface interactions at the ice/rock interface. Supported by data from the Southern Alps and other mountainous areas, I propose a simple numerical heat flow model to predict the climatic conditions and rock fracture spacing required for frost cracking. This model suggests that in highly fractured rocks the ultimate elevation of mountain peaks may be governed by the efficacy of frost action. This dissertation includes both my previously published and co-authored material.
vi
CURRICULUM VITAE
NAME OF AUTHOR: Tristram Hales
PLACE OF BIRTH: Christchurch, New Zealand
DATE OF BIRTH: 21 May 1979
GRADUATE AND UNDERGRADUATE SCHOOLS ATTENDED:
University of Oregon University of Canterbury
DEGREES AWARDED:
Doctor of Philosophy, 2006, University of Oregon Bachelor of Science with Honours (Geological Science), 2000, University of Canterbury
AREAS OF SPECIAL INTEREST:
Tectonic Geomorphology Landscape Evolution
PROFESSIONAL EXPERIENCE:
Graduate Teaching Fellow, University of Oregon, 2001-2006
Glacier Guide, Franz Josef Glacier Guides, 1999
vii
GRANTS, AWARDS AND HONORS:
PRIME Lab Seed Analysis Grant, Cosmogenic Isotope Sampling of Headwall Erosion, Purdue University, 2004
Graduate Student Research Award, Scree Production and Landscape Evolution, Geological Society of America, 2004
Outstanding Student Paper Award, Estimates of Erosion Rates for the Central Southern Alps, New Zealand, American Geophysical Union, 2002
PUBLICATIONS:
Hales, T.C., D.L. Abt, E.D. Humphreys, and J.J. Roering, (2005) Convective Instability as an origin for the Columbia River Basalts and Wallowa Mountain uplift in NE Oregon, USA. Nature, 438, 842-845.
Hales, T.C., and J.J. Roering (2005). Climate controlled variations in scree production, Southern Alps, New Zealand. Geology, 33, 701-704.
Hales, T.C., and J.J. Roering, (2006), Climatic controls on frost cracking and the evolution of bedrock landscapes. Journal of Geophysical Research. In press.
viii
ACKNOWLEDGMENTS
With the acceptance of this dissertation I will become the first person in my family to be awarded a Ph.D. The award is not the achievement of an individual, but instead is a testimonial to the love and support of my whanau, who are dearest to me. It has been my whanau that has taught me to love, appreciate and learn from nature, to work and play hard, and to enjoy life, and it gives me pleasure to dedicate this thesis to them.
I would like to say many thanks to Josh Roering, my advisor who has provided
me with much financial, academic, and emotional support over the last five years. Not
only is he an example of a great scientist, but an example of a balanced human being. He
is not only a great advisor and a great example of how to take the most from life. I really
appreciate his friendship and collaboration and hope that it continues into the future.
I am also grateful for the support of Gene Humphreys, who opened my horizons
beyond process geomorphology to thinking about the dynamics of the earth at a large
scale. His approach to science, particularly his focus on understanding, using simple
physics, the processes that govern the evolution of Earth’s surface have been an
inspiration.
To the rest of the faculty at the University of Oregon who have provided me with
help, support and friendship over the past five years. In particular thanks to Andrew
Marcus, Ray Weldon, Becky Dorsey, and Alan Rempel, who have provided great ix scientific support throughout my dissertation. I would also especially like to thank Kathy
Cashman and Marli Miller for exemplary teachers.
I would also like to recognize the support of the many friends that I have made during my time in Oregon, in particular, Kate Scharer, Todd Lamaskin, Heather Wright,
Reed Burgette, Noah Fay, and Darwin Villagomez for making the Geology department such a great place to work. Also many thanks to the people who kept my life balanced by taking me into the outdoors, watching world rally, teaching me about American sports, and eating Sunday brunch, particularly Pete and Sarah Erslev, Heather and Brian Ulrich,
Jake Casper, Garner Britt, Jacob Selander, and George (Mac) Jenkins. Finally to the many people that made my day to day life happy in the Roering lab, many thanks to
Alison Rust, Mackenzie Johnson, Molly Gerber, Suzanne Walther, Amanda Macleod,
Ben Mackey, and Lisa Emerson.
Last but by no means least this dissertation would not have happened without the love and support of my wife, field assistant and the best field camp student that I have ever had, Jenny Riker. x
For Ralph Edward Thomas Hales and Joie Louisa Lyttle
xi
TABLE OF CONTENTS Chapter Page I. INTRODUCTION ...... 1
II. A LITHOSPHERIC INSTABILITY ORIGIN FOR COLUMBIA RIVER FLOOD BASALTS AND WALLOWA MOUNTAINS UPLIFT IN NORTHEAST OREGON ...... 5
III. CLIMATE-CONTROLLED VARIATIONS IN SCREE PRODUCTION, SOUTHERN ALPS, NEW ZEALAND ...... 17 3.1. Introduction ...... 17 3.2. Study Area ...... 19 3.3. Methods ...... 21 3.4. Results ...... 23 3.5. Discussion and Conclusions ...... 25
IV. ROCKFALL EROSION AND POSTGLACIAL EVOLUTION OF THE SOUTHERN ALPS, NEW ZEALAND ...... 33 4.1. Introduction ...... 33 4.2. Study Area ...... 38 4.2.1. Western Southern Alps ...... 42 4.2.2. Axial Southern Alps ...... 44 4.2.3. Eastern Southern Alps ...... 45 4.2.4. Topographic Characteristics...... 50 4.3. Field and Topographic Analysis of the Eastern Southern Alps ...... 54 4.3.1. Statistical Analysis of the Elevation of Scree Slopes...... 54 4.3.1.1. Methods...... 54 4.3.1.2. Results...... 55 4.3.2. The Effects of Aspect ...... 56 4.3.2.1. Methods ...... 56 4.3.2.2. Results...... 59 4.3.3. Rock Fracture Spacing ...... 61 4.3.3.1. Methods ...... 61 4.3.3.2. Results...... 62 4.3.4. Headwall Erosion Rates ...... 64 4.3.4.1. Methods ...... 64 4.3.4.2. Results...... 65 4.4. Discussion ...... 67 4.4.1. The Effects of Glacial Cycles on Scree Production ...... 68 4.4.2. Landscape Development of Glacial-Interglacial Timescales...... 72 xii
Chapter Page 4.5. Conclusions ...... 73
V. CLIMATIC CONTROLS ON FROST CRACKING AND IMPLICATIONS FOR THE EVOLUTION OF BEDROCK LANDSCAPES ...... 75 5.1. Introduction ...... 75 5.2. Mechanisms for Rock Fracture by Ice Growth ...... 80 5.3. Model Setup...... 84 5.4. Model Assumptions...... 86 5.5. Model Results ...... 89 5.5.1. Timing of Segregation Ice Growth ...... 89 5.5.2. The Influence of Mean Annual Temperature ...... 92 5.5.3. The Influence of Annual Temperature Variations...... 96 5.5.4. Summary Effects of Ta and MAT...... 98 5.6. Discussion ...... 101 5.6.1. Segregation Ice Growth and Rock Fracture...... 102 5.7. Comparison to Previous Studies ...... 104 5.7.1. Timing of Rockfall Erosion...... 104 5.7.2. Southern Alps, New Zealand ...... 105 5.7.3. Utah Rockfall Inventory...... 108 5.7.4. Implications for Mountain Range Evolutions...... 111 5.8. Conclusions...... 113
BIBLIOGRAPHY ...... 115 xiii
LIST OF FIGURES
Figure Page 2-1. Regional structural and location maps for NE Oregon...... 8 2-2. Maps of post-eruptive uplift in the Wallowa Mountains area ...... 10 2-3. Interpolation of mapped lava flow distributions ...... 11 2-4. Tomographic images of NE Oregon...... 13 3-1. Map of the tectonic setting of New Zealand and distribution of scree slopes.....19 3-2. Graphs of topography, scree slopes, uplift rates and erosion rates...... 20 3-3. Annual air temperature variations in the Southern Alps, New Zealand ...... 30 4-1. Photographs of differences in the style of erosion across the Southern Alps...... 34 4-2. Estimates of rock uplift and erosion rates ...... 36 4-3. Map of the tectonic setting of New Zealand ...... 39 4-4. Distribution of scree slopes across the Southern Alps, New Zealand ...... 48 4-5. The distribution of elevation, slope, local relief and mean drainage density ...... 51 4-6. Plot of the mean scree slope and hillslope elevations ...... 55 4-7. Rose diagrams showing the aspects of scree slopes across the Southern Alps...57 4-8. Rose diagrams showing the aspects of hillslopes across the Southern Alps...... 58 4-9. Plots of aspect as a function of elevation...... 60 4-10. Block size and joint spacing data for the Cass area, Southern Alps...... 63 4-11. Plot of bedrock headwall and scree slope ages, Southern Alps...... 66 4-12. Shaded relief maps of the Cragieburn Range, Southern Alps...... 71 5-1. Oblique aerial photograph of the Cragieburn Range, Southern Alps...... 76 5-2. Results for our heat flow model at MAT of 0.001, Ta of 12oC...... 90 5-3. Results for our heat flow model at MAT of -5oC, Ta of 12oC...... 91 5-4. Changes in cracking intensity as a function of MAT for Ta of 12oC...... 93 5-5. The number of days spent in the frost cracking window...... 94 5-6. Changes in cracking intensity with depth...... 97 5-7. Summary plots of the effects of changes in MAT and Ta ...... 100 5-8. Yearly atmospheric temperature variation in the Southern Alps...... 107 5-9. Elevation and frequency of rockfall events on roads in Utah...... 109 xiv
LIST OF TABLES
Table Page 3-1. Calculated erosion rates for debris and alluvial fans in the Southern Alps...... 25 3-2. Rock quality designation (RQD) measurements...... 26
1
CHAPTER I
INTRODUCTION
The idea that landscapes evolve due to the interaction of uplift forces and erosional processes has a long history. Inspired by the work of Charles Darwin
[1859], William Morris Davis [1899] suggested that mountains evolve in similar way to plants and animals. He suggested that every landscape develops from a youthful state, usually an uplifted plateau with few incised streams, through maturity (greater dissection) to old age, where a peneplain forms. The fundamental flaw of a descriptive classification such as this one is that landscapes do not follow a simple path of uplift and degradation; instead the inputs are constantly changing, with uplift rates varying with changes in tectonics and erosion. Davis’ model dominated the geomorphic literature until the 1960s, when rediscovery of the work of Grove Karl
Gilbert, led to a change in research direction towards an understanding of the processes that govern how mountains uplift and erode.
Students of Gilbert view the evolution of mountainous landscapes as an interaction between uplift processes, including horizontally-directed plate tectonic forces and vertically directed isostatic uplift, and erosional processes, such as fluvial, glacial, periglacial, and hillslope erosion [Adams, 1980; Hack, 1960; Reneau and
Dietrich, 1991; Willett and Brandon, 2002]. Uplift and erosion are strongly coupled; 2
increasing the uplift rate in a mountain range will steepen slopes and strengthen erosion rates [Gilbert, 1877]. It is this negative feedback that suggests mountain ranges tend towards a steady state, where the amount of material entering a mountain range through tectonic processes is balanced by the amount that is removed by erosion [Hack, 1960; Willett et al., 2001]. Recently, geodynamic models of mountain belts have proposed that, on a million year timescale, mountains can reach a steady state topography [Brandon et al., 1998; Whipple and Meade, 2004; Willett et al.,
2001].
Mountain evolution is also modulated by climatic changes, which can change the style of erosion by creating glaciers or enhancing precipitation [e.g. Bull and
Knuepfer, 1987]. These climate changes are, arguably [see Raymo and Ruddiman,
1992], driven by external forcing such as changes in the Earth’s orbit and act on a tens to hundreds of thousand year timescales [Wunsch, 2004]. Changes in climate, particularly changes from fluvial to glacial erosion cause isostatic uplift of mountain peaks [Molnar and England, 1990]. While the timescale of tectonic forcing and climate change differ by (at least) an order of magnitude, how their interaction affects the overall shape and development of orogens with time is an important question in geomorphology today.
New Zealand’s Southern Alps have been an important test case for the concept of steady state, because of high uplift and erosion rates and because the orogen can be modeled with a numerically simple wedge-type of model [Batt and 3
Braun, 1999; Beaumont et al., 1996; Koons, 1990, 1994; Willett et al., 2001]. The
Southern Alps is formed by an oblique continental collision in which the Pacific plate is being thrust over the Australian plate along the Alpine Fault. Uplift rates decay exponentially from 12 mm/yr in the west (at the Alpine Fault) to < 1 mm/yr in the east [Norris and Cooper, 2000; Tippett and Kamp, 1993; Wellman, 1979]. A strong orographic precipitation gradient exists across the Southern Alps, with precipitation rates of up to 15 m/yr in the west decaying to 1m/yr in the east. The strong uplift and precipitation gradients cause dramatic changes in the style of erosion across the range. Three erosional regions exist; a western region of fluvial dissection and landslides, an axial region of glacial erosion, and an eastern region of braided rivers and scree slopes [Adams, 1980; Whitehouse, 1988]. Erosion in the western part of the mountain range is well characterized with high, landslide-driven erosion rates of up to
18 mm/yr measured from sediment yields in streams [Griffiths, 1979; Griffiths, 1981] and from landslide inventories [Hovius et al., 1997; Korup, 2005c]. Erosion across the rest of the range has not been characterized in detail, despite evidence for efficient glacial erosion [Hicks et al., 1990] and rockfall, which produces kilometer-scale scree slopes [Whitehouse and McSaveney, 1983]. Chapters III and IV characterize the dominant controls on erosion in the eastern Southern Alps and the rates at which they occur by presenting field-data, aerial photograph mapping, cosmogenic radionuclide age dates, and estimates of fan volumes. These chapters show that erosion in the eastern Southern Alps is dominated by rockfalls caused by segregation ice weathering 4
of bedrock. Using these data I have formulated a numerical heat flow model that predicts the conditions at which ice is likely to break rocks and create rockfall
(Chapter V).
Inherent in most models describing the evolution of mountainous landscapes is that concept that erosion is driven by horizontally-directed plate tectonic forces
[e.g. Willett and Brandon, 2002]. However, recent work in the southern Sierra
Nevada Mountains has shown that vertically-directed isostatic forces, driven by the delamination of lithosphere can also produce rapid uplift [Ducea and Saleeby, 1996;
Saleeby and Foster, 2004]. In northeastern Oregon, rapid uplift of the Wallowa
Mountains is coincident with eruption of the Columbia River Basalt Group, a thick accumulation of flood basalt lavas. Chapter II describes an effort to quantify the uplift of the largest granitic pluton in an area of tectonic quiescence and suggests that uplift may have been driven by delamination of lithosphere from the base of the granitic
Wallowa pluton. The coincidence of Wallowa Mountains uplift and Columbia River
Basalt volcanism suggests a link between them, such that delamination of granitic roots may allow decompression melting of asthenospheric mantle, causing flood basalt volcanism. 5
CHAPTER II
A LITHOSPHERIC INSTABILITY ORIGIN FOR COLUMBIA RIVER FLOOD
BASALTS AND WALLOWA MOUNTAINS UPLIFT IN NORTHEAST OREGON
This chapter was published in the journal Nature in 2005, volume 438,
pages 842-845. This paper was coauthored with David Abt, who
conducted the tomographic analysis and provided a figure, Gene
Humphreys and Josh Roering, who provided funding and editorial
help.
Flood basalts often initiate hotspot magmatism. The Columbia River Basalts
(CRB) volumetrically dominate flows initiating the Yellowstone hotspot, yet their source is ~500 km north of the projected hotspot track, in the vicinity of the Wallowa
Mountains [Camp and Ross, 2004]. These mountains are composed of the largest granitic pluton in an area characterized by lithosphere of oceanic affinity [Lallemant,
1994]. Mapped CRB flow-interface elevations indicate mild pre-eruptive subsidence of the Wallowa Mountains, followed by syn-eruptive uplift of several hundred meters and a persistent uplift of ~2 km. Surface uplift mimics regional topography, which 6
has the Wallowa Mountains surrounded by a set of valleys which themselves are surrounded by a circular trend of lower mountains. Here we present tomographic mantle images that revealan area of reduced mantle melt content coincident with the
200-km wide topographic bull’s eye. Convective downwelling and detachment of compositionally dense plutonic roots can explain the timing and magnitude of CRB magmatism and surface uplift and the location of the imaged melt-depleted mantle.
The Yellowstone hotspot initiated ~17 Ma with a rapid succession of large eruptions starting at McDermitt caldera in northern Nevada and propagating ~500 km north along the western edge of the Precambrian continental margin, first to Steens
Mountains area in SE Oregon and then to the site of the CRB eruptions in NE Oregon
[Camp and Ross, 2004] (Fig. 2-1). The total volume of erupted flood basalt is
~234,000 km3, 75% of which are CRB [Camp et al., 2003]. The large geographic extent involved in these eruptions has been difficult to reconcile with a traditional plume model because nearly all flood basalt erupted north of the projected hotspot track, which trends from Yellowstone to McDermitt caldera. Furthermore, the Snake
River Plain is a well-developed hotspot track, but it leads away from the CRB source area and arcs across the state of Idaho like a mischievous grin (Fig. 2-1). Two recent suggestions accounting for the distributed and largely off track nature of hotspot initiation modify the plume hypothesis by having plume impingement beneath either the Steens Mountains [Camp and Ross, 2004] or the central Idaho craton [Jordan,
2005] and then flattening rapidly along a trend near the Precambrian continental 7
margin. These plume models are supported by the presence of high 3He/4He in the flood basalts [Dodson et al., 1997]. In contrast, other models suggest that flood basalt magmatism represents passive back arc processes [Anderson, 1994a].
Currently little deformation occurs in NE Oregon. GPS and geologic data show clockwise rotation of the intact Oregon Coast block relative to North America about a pole near the corner of the 0.706 line (Fig. 2-1), consistent with regional tectonics [McCaffrey et al., 2000; Wells et al., 1998]. E-W extension on normal faults occurs with increasing rate south of this location, and N-S contraction, expressed as
E-W oriented folds, is increasingly active to the west. This regional tectonic setting has persisted since eruption of the CRBG without noticeable change [Hooper, 1997;
Wells et al., 1998], and proximity to the pole implies very low deformation rates for
NE Oregon. The regional ~N-S axis of maximum compression [Camp and Hooper,
1981; Hooper, 1997; Hooper and Camp, 1981; Hooper and Conrey, 1989] responsible for these faults and folds also was responsible for the NNW-SSE trending
Chief Joseph dike swarm, E-W folds such as the Lewiston anticline and NW-SE oriented strike-slip structures such as the Olympic-Wallowa Lineament [Camp and
Hooper, 1981; Hooper and Camp, 1981; Hooper and Conrey, 1989]. Young deformation near the Wallowa Mountains, however, is exceptional and inconsistent with deformation associated with the regional strain field. Also there are a number of structures with orientations that are inconsistent with a ~N-S regional strain field including the oblique normal Limekiln and Hite Fault systems, the NE-SW oriented 8
-120 -116 -112 WA 48
Precambrian Continental Margin Y 44 S SRP M
B OWL
Elevation W 3km
2
1
OR ID 0
Figure 2-1. Regional structural and location maps for NE Oregon. (a) Map of western United States showing locations of major magmatic centres related to the Yellowstone hotspot (Y), including: McDermitt caldera (M), Steens Mountains (S), the arcuate Snake River Plain (SRP), and the source of the CRB eruptions (C). Note the bull’s eye pattern of uplift at the NW end of the SRP. (b) Location map showing CRB source area, the Wallowa (W) and Blue (B) Mountains. Productive Grande Ronde dikes are shown as short lines drawn at an average orientation [Camp and Ross, 2004]. Major extensional (lines with balls) and contractional (double arrows) structures are indicated [Hooper and Conrey, 1989]. The Precambrian continental margin, represented with 87Sr/86Sr=0.706, is indicated with a dashed line. Major plutons west of Precambrian North America are outlined with heavy dotted lines. 9
folds such as the Troy Basin syncline [Hooper and Conrey, 1989]. We suggest that these inconsistencies in interpretation represent local perturbations to the regional strain field due to the imposition of a local, vertically-directed strain.
CRB lavas erupted onto a low relief surface of Mesozoic accreted ocean terranes stitched together by Jurassic plutons, of which the Wallowa pluton is the largest [Goles, 1986] (Fig. 2-1). After initial Imnaha flows filled the valleys, subsequent eruptions deposited thin, flat-lying sheets that provide a useful datum to measure deformation accurately. These flows are found in Hells Canyon at an elevation of 1100 m and atop the Wallowa Mountains at 2700 m, demonstrating substantial vertical motion following their eruption [Walker, 1979]. Grande Ronde eruptions followed the Imnaha, depositing about 80% of the CRB lavas over a duration of ~1 m.y. We correct elevation maps of the Grande Ronde magnetostratigraphic units [Hooper et al., 1985; Reidel et al., 1992; Walker, 1979] for the effects of erosional unloading[Anderson, 1994b; Lambeck, 1988] to obtain markers of post-eruptive deformation (Fig. 2-2 & 2-3). The uplift and downwarp of these flows closely conform to current topography, with up to 2 km of uplift focused on the Wallowa pluton (Figs. 2-1 & 2-2) and more subdued uplift centred on the lesser plutons in the area (Fig. 2-1). The history of Imnaha and Grande Ronde uplift reveals the early evolution of local buoyancy. Imnaha basalts ponded in incipient basins, indicating possible pre-eruptive subsidence [Camp, 1981]. Basin development continued during Grande Ronde eruptions, causing local ponding and 10
Figure 2-2. Maps of post-eruptive uplift in the Wallowa Mountains area. (a) Digital elevation map showing the overlapping distribution of exposed Grande Ronde magnetostratigraphic interfaces used for analysis of Wallowa Mountains uplift: Imnaha-R1 (blue), R1-N1 (green), N1-R2 (gold) and R2-N2 (red). (b) Composite surface for Grande Ronde uplift. This surface was made by vertically shifting the individual interfaces using published flow thicknesses [Walker, 1979], smoothing the resulting surface with a mild low-pass filter, and removing the effects of erosional unloading by deconvolving the point load response of an elastic plate [Anderson, 1994b; Lambeck, 1988], assuming total coverage by the Grande Ronde lavas and an elastic thickness of 5 km similar to other estimates from this area [Saltzer and Humphreys, 1997]. The absolute elevation is referenced to a hinge line (black line) that separates the uplifted Blue Mountains from the down-dropped Pasco Basin immediately NW of the Blue Mountains. (c) Schematic plot of the Wallowa Mountains uplift history (dark line) and a hypothetical area affected by a thermal plume (grey line). Wallowa Mountain uplift history is controlled at three times. Pre-eruptive subsidence, inferred from ponding of early Imnaha lava flows in basins, is interpreted as the initiation of delamination. We calculated ~300 m of syn-eruptive uplift in this area over the 1 m.y. duration of Grande Ronde eruption. This is followed by rapid uplift of the Wallowa Mountains 3-4 m.y. after initial Imnaha eruption, evidenced by the development of structures bounding the Wallowa Mountains [Walker, 1979], and current topographic relief of the composite surface shown in (b). Note that the hypothetical uplift curve for a thermal plume can vary significantly in magnitude, we have chosen a representative amount of uplift predicted CRB magmatism. 11
thickening of flows, while syn-eruptive uplift created locally thinned units. In particular, ~300m of uplift resulted in a thinned N2 unit near its source close to the
Wallowa Mountains [Camp, 1981; Reidel et al., 1989] (Fig. 2-3). Post-Grande Ronde deformation is evidenced by the contrasting distributions of Grande Ronde and late
CRB Wanapum (15-14 Ma) and Saddle Mountains (14-6 Ma) flows, which show distinctive, well-developed channelization and ponding in response to emerging topographic relief [Tolan et al., 1989]. Significant relief generation occurred at 14.5-
51 51 Saddle Mountains
) Tsm 5 0 5 1
x (
g n i h
rt Tw Wanapum o N Northing (x 10 (x Northing ) Tgn2
50 Tgr2 Grande Ronde Tgn1 50 4550 5 55 Easting (x 10 ) Tgr1 5 45 Easting (x 10 ) 50
Ti Imnaha
51
2.5
5 51
2.0 5
1.5 Northing (x 10 )
1.0 10 (x Northing ) Elevation (km) Elevation
50 0.5
0 50
40 45 50 55 45 50 55 Easting (x 105 ) Easting (x 105 )
Figure 2-3. Interpolation of mapped lava flow distributions. Digital elevation models showing the elevation of each surface interpolated from the current mapped extent of flow interfaces (black dots). The overlapping extent of each flow was then using areas that were well constrained with data and published thicknesses to create a composite surface (Fig 2-2). Dashed lines show the locations of cross sections. 12
13 Ma coincident with initiation of the La Grande, Baker, and Weiser grabens, which surround the Wallowas Mountains [Camp, 1981]. This structural evidence leads us to conclude that most Wallowa Mountains uplift occurred in <10 m.y. (Fig. 2-2c).
The current physical state of the upper mantle provides constraint on the origin of buoyancy. We obtain a basic assessment of upper mantle structure by tomographically inverting ~600 teleseismic P-wave delays collected by a six-station array of IRIS-PASSCAL seismometers deployed in the region. Relatively high- velocity mantle is imaged beneath the bull’s eye region (Fig. 2-4). Resolution
“squeezing” tests [Saltzer and Humphreys, 1997] (Fig. 2-4b) indicate that this seismically fast volume extends from ~75 km to at least at least 175 km depth, and its horizontal extent is coincident with the bull’s eye pattern (Fig. 2-4). The occurrence of basaltic volcanism in the recent past and the major uplift in the Wallowa
Mountains are inconsistent with the high-velocity mantle being relatively cool. The only reasonable alternative is that melt content has been reduced (and possibly eliminated) by eruption of the CRB from the high velocity volume
[Hammond and Humphreys, 2000; Saltzer and Humphreys, 1997]. This hypothesis is strengthened by the spatial correlation between the upper mantle high-velocity volume and the location of flood basalt dykes (Figs. 2-1 & 2-4). The implied mantle depletion would increase mantle buoyancy by an amount that accounts for the observed several hundred metres of syn-eruptive uplift [Schutt and Lesher, 2005]. A back of the envelope calculation suggests that 5% depletion of an 80-km thick layer 13
(e.g. 70-150 km) would cause ~300m of isostatic uplift in the region. However mantle basalt depletion cannot account for the focused ~2km of surface uplift of the
Wallowa Mountains.
Figure 2-4. Tomographic images of NE Oregon. (a) Upper mantle P-wave velocity variations at 100 km depth, imaged by teleseismic tomography from data recorded by six seismometers (triangles). Also shown are: the location of state boundaries, the Wallowa pluton (cross- hatched area) and cross section location (dashed line). (b) Cross-sections through squeezed inversions. These are hypothesis tests in which the structure that best accounts for the observed delays is found that is confined above the green line; after relaxing the depth constraint, creation of significant structure below the green line indicates that deeper structure is required. The upper cross section (squeezing depth=100 km) indicates a strong need for structure below 100 km. The middle panel (squeezing depth=150 km, and used for (a)), like the lower panel (unconstrained) indicates no need for high-velocity structure below 150 km depth. Imaging places the top of the high-velocity body near 75 km.
The unusual tectonic and magmatic history of NE Oregon has similarities to two recently-studied events that involve apparent small-scale convective instability driven by dense garnet-rich crust and mantle lithosphere. Elkins-Tanton and Hager 14
[2000] account for pre-eruptive subsidence of the Siberian traps flood basalts with the initial downwelling of dense eclogitic lower lithosphere beneath the area of downwarping, and attribute the magmatism to decompression melting of ascending asthenosphere flowing into the vacated regions. Saleeby and Foster [2004] compile abundant evidence for young and ongoing convective descent (termed “delamination” in their paper) of the garnet-rich rocks that are the genetic compliment to the granitic
Sierra Nevada pluton in California. As this dense lower crust and upper mantle sinks, the overlying southern Sierra Nevada Mountains rise in isostatic response.
In NE Oregon, where granitic plutons are scarce, major uplift is concentrated on the largest pluton in the region. We attribute this uplift to the convective removal of the pluton’s dense roots, and consider the magnitude of uplift and its strong association with the Wallowa pluton difficult to account for with any other process.
Furthermore, the remarkable bull’s eye uplift pattern, the post-eruptive timing of most uplift, and the associated CRB outpourings, can be explained with such a delamination-driven model. We envision initial convective instability of the plutonic roots creating a drip-like downwelling that pulls the Wallowa region down, creating pre-eruptive subsidence (Fig. 2-2c). Asthenosphere flowing up into the area around the Wallowa downwelling decompresses and partly melts upon ascent, creating initial
CRB flows that pond in the depression. As the dense mantle and lower crust sink and detach from the lithosphere, surface uplift begins and magmatism accelerates within the circular area where decompression melting is excited. The uplift of ~300 m that 15
occurred during this time interval is controlled by the combined effects of melt accumulation in magma chambers and chamber evacuation, and residuum buoyancy created by basalt depletion. Finally, continued sinking of the dense Wallowa pluton root causes continued uplift of the Wallowa pluton, which progresses until the sinking mass is effectively removed from the region and uplift becomes fully isostatic (Fig. 2-
2c).
This model for flood basalt volcanism driven by delamination of compositionally dense roots has some interesting implications for Yellowstone hotspot initiation. First, if a plume impacted to the south of the CRB source area and propagated north along the Precambrian continental margin [Camp and Ross, 2004;
Jordan, 2005] associated hot mantle material could trigger delamination of the
Wallowa plutonic roots. This mechanism could therefore provide a way of accounting for the intense magmatism far from the Yellowstone hotspot track. In this case, a significant modification to the plume hypothesis is not required and the problem of
CRB magmatism is addressed. A second interpretation is that the sequence of eruptions initiating the Yellowstone hotspot was caused by a series of delamination events, perhaps one triggering the next or all triggered by unusual back-arc activity, and with no requirement for an active plume. In this case, magmatism could be related to a swath of fertile mantle lithosphere adjacent to the infertile Precambrian lithosphere that was recently hydrated in a fore-arc setting, thereby integrating into the magmatic history this peculiar association with the sharply truncated Precambrian 16
margin of North America. In either case, delamination is required to initiate flood basalt volcanism in NE Oregon, and an intimate relationship between crustal structure and flood basalt initiation is demonstrated.
This work inspired by conversations with the late Dr. Gordon Goles. Funding for this work provided by National Science Foundation grants EAR-0309975 to JJR and EAR-0106892 to EDH. Reviews from Mark Richards and two anonymous reviewers greatly improved the manuscript. We would like to thank Jenny Riker,
Dawn Clippinger, Cody Simmons and Marc Giba for field assistance. We would also like to thank Noah Fay and Jason Crosswhite for help with numerical analysis 17
CHAPTER III
CLIMATE-CONTROLLED VARIATIONS IN SCREE PRODUCTION,
SOUTHERN ALPS, NEW ZEALAND
This chapter was published in the journal Geology in 2005, volume 33
(9), pages 701-704. The paper was co-authored with Josh Roering,
who provided advisorial and editorial support and funding for the
project.
3.1. Introduction
Bedrock erosion by hillslope processes such as bedrock landsliding, rock avalanches, and debris flows significantly affects the size and shape of mountains
[Burbank et al., 1996; Schmidt and Montgomery, 1995], and rates of sediment production by bedrock landslides and debris flows are often significant relative to rock uplift and fluvial incision [e.g. Hovius et al., 1997; Reneau and Dietrich, 1991].
In mountainous areas, scree-mantled slopes are ubiquitous, and localized studies of headwall retreat show that rockfall occurs at geomorphically significant rates (up to
4.5 mm/yr) [see reviews in Ballantyne and Eckford, 1984; Sass and Wollny, 2001]. 18
However, the contribution of small-scale (<102 m3) rockfalls is seldom addressed in regional assessments of mountain erosion, because such small events are difficult to document across large areas.
The regional pattern of rockfall and scree production reflects the integrated effects of processes typically studied at the scale of individual headwalls. Controls on rockfall activity in alpine environments include availability of erodable bedrock
[Brook and Tippett, 2002], time spent in a periglacial environment, freeze-thaw cycles, frost cracking, frost wedging [Matsuoka, 2001], vegetation (which both adds cohesive strength to bedrock through root systems and reduces rock strength via chemical and mechanical weathering), glacial loading (which creates oversteepened headwalls and fractures rock) [Miller and Dunne, 1996], and seismic activity
[Matsuoka, 2001].
Although it has been proposed that fundamental morphologic characteristics of mountain ranges, such as relief, slope, and drainage density, are controlled by fluvial and glacial incision [Whipple et al., 1999] and large landslides [Schmidt and
Montgomery, 1995; Wieczorek, 2002], the topographic imprint of scree-generating processes is poorly constrained. In contrast to intensive, field-based studies of headwall dynamics, we use the spatially and temporally averaged signature of rockfall activity to quantify the contribution of scree production across the Southern
Alps of New Zealand. The distribution of scree slopes reflects the accumulation of rockfall events, and debris and alluvial fans emanating from scree-dominated basins 19
provide a postglacial record of scree production. Systematic variations in rock uplift rate, precipitation, and temperature across our study area enable us to relate scree production to climatic- and tectonic-driven factors.
3.2. Study Area
The Southern Alps are formed by oblique collision of the Pacific and
Australian plates along the Alpine fault [Walcott, 1998] (Fig. 3-1). Uplift rate decays from ~11 mm/yr at the Alpine fault to <1mm/yr at the eastern range front [Adams,
Cass 168o E N o Otira 40 S
lt au F ne pi Al
Lake Coleridge
0 20 40 60 80 Distance from the Alpine Fault (km)
Figure 3-1. Map of the tectonic setting of New Zealand and distribution of scree slopes. Subduction occurs in the north and south margins, and continental collision occurs along the Alpine fault. Inset is a shaded relief map of the transect across the Southern Alps analyzed in this study. Superimposed on the map is a location of vegetated (red) and unvegetated (orange) scree slopes estimated from aerial photo mapping. Values at the bottom of the inset map represent distances (km) from the Alpine fault.
20
1980] (Fig 3-2). A prevailing westerly airflow creates a strong orographic effect with maximum precipitation mid-slope on the western range front of 15 m/yr, decreasing
2500 2500 MAP Main Divide A 12 C Max. Elev. Mean ± σ 2000 2000 8 1500 1500 Mean Elev. 1000 1000 4 MAP (m/yr) Elevation (m) Elevation 500 Min. Elev. 500 Active Scree Elevation (m) Elevation Scree Active 0 0 0 0 20406080 0 20406080 D 80 B 100 Landslide Total scree Rock 60 (Active + Veg.) 10 Uplift
40 1 Susp. Sed.
20 0.1 Scree/Alluvial Fans Active scree Scree (% of hillslope area) hillslope of (% Scree 0 Rate (mm/yr) Erosion or Uplift 0.01 0 20406080 0 20406080 Distance from Alpine Fault (km) Distance from Alpine Fault (km)
Figure 3-2. Graphs of topography, scree slopes, uplift rates and erosion rates. A: Graph showing the minimum (light gray), mean (dark gray), and maximum (black) elevations calculated for 2-km swaths across the transect. Black squares represent mean annual precipitation (MAP) [Griffiths and McSaveney, 1983a]. Note the strong orographic effect from west to east across the range. B: Minimum (light gray squares) and maximum (dark gray squares) estimates of the fraction of hillslopes mantled by scree across our transect. Note the rapid increase in scree mantling in the eastern part of the range. C: Variation in the elevation of active scree deposits (orange zones in Fig. 3-1). Solid black line and shaded region represent the mean and standard deviation of elevation values. Note the relatively consistent elevation of the scree slopes. D: Distribution of rock uplift and erosion rate estimates. Uplift estimates, using geologic indicators and tilt of paleolake terraces [black squares; Adams, 1980], are approximately equal to estimates of scree production in the eastern region [black triangles; this study] and bedrock landsliding in the west [light gray circles; Hovius et al., 1997]. Estimates of bedrock erosion from suspended sediment [dark gray circles; Griffiths, 1981] are generally lower than for scree production in the eastern part of the Southern Alps. 21
to <1 m/yr in the east [Griffiths and McSaveney, 1983a] (Fig. 3-2A). The Southern
Alps have been extensively glaciated, and during the last glacial maximum (LGM), glaciers advanced beyond the range front on the northwest and southeast sides of the orogen [Suggate, 1990]. In the west, U-shaped valleys have been fluvially dissected during the current interglacial, generating steep, vegetated slopes prone to bedrock landsliding. East of the Main Divide, broad glacial valleys have been infilled by active braided streams and debris and alluvial fans [Adams, 1980]. Currently, glaciers are found only at relatively high elevations along the Main Divide. West of the Main
Divide, erosion rates via bedrock landsliding and suspended sediment analyses yield estimates of ~2–18 mm/yr and ~6–12 mm/yr, respectively, [Griffiths, 1981; Hovius et al., 1997]. In contrast, erosion rates in the eastern region of our transect have values
<1 mm/yr [Griffiths, 1981] (Fig. 3-2D).
3.3. Methods
The spatial distribution of scree deposits in the Southern Alps was quantified by aerial photograph mapping within a 40 by 80 km transect (Fig. 3-1). Our study transect is bounded by the Alpine fault at the western margin and the Canterbury
Plains to the east. We generated two estimates for the extent of scree-mantled slopes; the first represents areas covered by active scree, and the second includes paleoscree slopes that are currently mantled by vegetation. These methods provide minimum and maximum estimates of scree distribution in our transect, respectively. Active scree 22
was characterized as areas of light color on the aerial photo that showed a slope- parallel fabric, had a source bedrock outcrop, and exhibited consistently moderate
(~30°) slope angles (determined from digital elevation model analysis). We identified areas of vegetated scree as having similarly moderate slopes combined with evidence for exposed scree where vegetation has been locally removed. Because our analysis focuses on hillslopes, we removed valleys in our calculation of the proportion of slopes mantled by scree. To characterize how the distribution of scree varies with tectonic and climatic variables, we subdivided our transect into 40 2-by-40-km swaths oriented with the long axis parallel to the Alpine Fault.
We calculated long-term rockfall erosion rates in the eastern part of our transect by estimating the volume of debris and alluvial fans accumulated on surfaces of known age. We analyzed fans formed at the mouths of small (<1 km2), unglaciated drainage basins, where field observations indicated that rockfall activity dominates sediment production. During the LGM these small, steep drainage basins drained directly onto large valley glaciers, such that most sediment produced was likely evacuated. The fans that developed within the Cass Valley and Lake Coleridge sites accumulated after retreat of large valley glaciers at ca. 13.5 ka [Suggate, 1990] (Fig.
3-1). The Cass Valley was isolated from further glacial and fluvial erosion by a bedrock ridge at the head of the valley such that the fans provide a complete record of post-LGM sediment production from the small basins. In the western half of our transect (~20 km from the Alpine fault), we estimated rockfall erosion rates using 23
scree deposits situated along the margin of valley walls that were occupied by glaciers at 10–12 ka [Ivy-Ochs et al., 1999].
For the eastern sites, we estimated the volume of each debris or alluvial fan using a 25 m digital elevation model, interpolating proximal valley wall and floor elevations as basal datums. In the western sites, we used ground-based surveys to estimate the area of scree deposits and applied a caternary curve to valley profiles to estimate the geometry of underlying bedrock slopes [Hirano and Aniya, 1988]. Scree production rate was calculated by dividing the volume of each deposit by the time since deglaciation and basin source area. We generated bedrock erosion rate estimates by adjusting for the density of fan or scree material (1600 ± 100 kg m–3) and bedrock
(2700 ± 100 kg m–3) [Sass and Wollny, 2001]. Because the debris and alluvial fan calculations do not account for scree stored with basins above each fan, erosion rate estimates for the eastern sites reflect minimum values.
3.4. Results
The distribution of scree, which reflects rates of production and removal by fluvial processes, varies systematically across our transect (Fig. 3-2B). In the western
Southern Alps, scree deposits are small, sparse, and account for <10% of hillslope area (Fig. 3-1). These deposits are typically found beneath isolated, unvegetated headwalls at high elevations. Sparse evidence for relict (vegetated) scree slopes exists west of the Main Divide. The fraction of scree increases to 10%–20% at higher 24
elevations along the range axis, where scree deposits form along the margins of U- shaped valleys and are disconnected from the channel network, suggesting that the preservation potential is relatively high. Bedrock exposure is common in this area because much of the landscape is above the tree limit. East of the Main Divide the fraction of total (vegetated and active) scree-mantled slopes increases monotonically
(see dark squares, Fig 3-2B), reaching a maximum value (70%) between 40 and 65 km from the Alpine fault. The proportion of active scree remains relatively constant between 20% and 30% (see light gray squares, Fig. 3-2B). In this region, scree often mantles entire hillslopes, and some exceed 1 km in length. Convex, scree mantled ridges and small, isolated bedrock headwalls dominate the headwaters of small basins. Field observations suggest that scree production in this area is transport limited, such that active transport is required to expose bedrock faces.
To analyze how climatic variables may affect rockfall initiation, we calculated the elevation distribution of scree slopes across our transect. For this analysis, we used the distribution of active scree as it allows for comparison with the modern climatic regime. Despite systematic variations in precipitation, rock uplift, and relief, the mean elevation of active scree slopes (Fig. 3-2C) is surprisingly consistent at
~1400 m, and nearly 70% of the deposits occur between 1200 and 1600 m a.s.l.
Minimum erosion rates calculated from measurements of 15 debris and alluvial fans in the eastern region (Cass Valley and Lake Coleridge) have a mean value of 0.6 mm/yr with a standard deviation of 0.4 mm/yr (Table 3-1). Rates at three 25
of the sites exceed 1 mm/yr, and basins with high scree production rates appear to contain large stores of scree that inundate underlying ridge and valley topography.
Near the Main Divide, our field-based estimates of scree production rates are uniformly low (0.01 mm/yr).
Easting Northing Volume (m3) Age (a) Basin Area (m2) ER (mm/yr) 2411157 5788087 33200000 13500 2790000 0.58 2410180 5788087 22900000 13500 2350000 0.47 2412856 5784659 22400000 13500 4830000 0.22 2412435 5784088 6080000 13500 3590000 0.08 2413953 5789967 23700000 13500 1670000 0.69 2415968 5785080 6650000 13500 494000 0.65 2416314 5786268 30200000 13500 2920000 0.50 2410165 5789425 24400000 13500 891000 1.33 2414690 5789395 5680000 13500 808000 0.34 2415366 5789049 1870000 13500 322000 0.28 2409623 5795785 19200000 13500 1020000 0.92 2408781 5791981 4020000 13500 791000 0.25 2390607 5771620 25600000 13500 1150000 1.08 2388721 5773549 16700000 13500 692000 1.17 2389554 5768595 11500000 13500 1390000 0.40 2392579 5776443 14800000 13500 860000 0.83 2384074 5770568 16900000 13500 3760000 0.22
Table 3-1. Calculated erosion rates (ER) for debris and alluvial fans in the Southern Alps.
3.5. Discussion and Conclusions
Previous studies of individual headwalls have recognized many factors that control scree production, including rock type, vegetation, temperature, earthquakes, and time since deglaciation [e.g. Ballantyne and Eckford, 1984; Sass and Wollny,
2001], although the relative importance of these factors remains elusive. Our method uses regional patterns in scree distribution, which increases systematically from west to east, to infer how climatic and tectonic factors influence rockfall activity across the 26
Southern Alps. First, variations in lithology and metamorphic grade do not account for the observed scree distribution. Jointed sandstones and argillites of the Torlesse
Supergroup are dominant from 20 to 80 km along our study transect. These rocks have been deformed in at least four major regional events, such that no preferential trends in sandstone/shale ratio or joint orientation exist [MacKinnon, 1983]. As a result, systematic changes in mechanical rockfall susceptibility appear unlikely, as estimates of both rock mass strength [Augustinus, 1995a] and rock quality designation
[Deere and Deere, 1988] (Table 3-2) generate consistent values near the Main Divide and in the east region of our transect, despite strong variations in scree mantling (Fig.
3-2B). At 20 km along our transect, a discrete change occurs from sandstone-argillite to schist [MacKinnon, 1983]. Although it has been suggested that schists and other higher grade metamorphic rocks west of the Main Divide may obviate scree production [Whitehouse, 1988], we observe continuity in the distribution of scree- mantled slopes across this distinct lithologic boundary.
Over geologic time scales, large-magnitude earthquake events have frequently occurred along the Alpine fault, whereas seismic activity has been less persistent in the eastern part of our transect [Pettinga et al., 2001]. During large-magnitude earthquakes, rockfall activity is concentrated near the epicenter and decays with distance [Adams, 1980]. Given that the Alpine fault is the dominant seismic source in our study area, our observed scree distribution is inconsistent with earthquakes as the primary scree generating mechanism. 27
Easting Northing Elevation Distance from AF (km) RQD 2390103 5811616 1285 19.00 0.275 2389904 5811447 1275 19.01 0.468 2390395 5811655 1239 19.14 0.598 2390667 5811542 1080 19.39 0.623 2390746 5811577 1137 19.41 0.300 2407835 5791570 1219 45.68 0.333 2407630 5791331 1382 45.74 0.271 2408336 5791588 758 45.96 0.050 2408336 5791588 839 45.96 0.463 2409150 5789943 1239 47.77 0.432 2409003 5789804 1299 47.79 0.357 2409233 5789612 1112 48.08 0.055 2409333 5789477 1014 48.25 0.518 2408339 5788721 1631 48.25 0.413 2409460 5789558 932 48.26 0.624 2408920 5788527 1329 48.76 0.427 2409005 5788367 1203 48.94 0.473 2409585 5788799 953 48.94 0.138 2408564 5787700 1576 49.21 0.338 2409269 5787952 1043 49.43 0.570 2413148 5788883 1291 51.02 0.359 2413198 5788220 1588 51.58 0.528 2413462 5788206 1601 51.75 0.622 2413366 5788089 1674 51.78 0.630 2414068 5788586 1238 51.81 0.427 2414409 5788836 944 51.82 0.530 2414140 5788629 1161 51.82 0.493 2413351 5788024 1672 51.83 0.665 2412468 5787285 1357 51.89 0.540 2413553 5787822 1607 52.11 0.535 2414032 5788143 1365 52.14 0.237 2413283 5787508 1641 52.20 0.605 2413706 5787382 1629 52.55 0.593
Table 3-2. Rock quality designation (RQD) measurements. These measurements were made in Torlesse terrane rocks in the Southern Alps, New Zealand.
Glaciation has been proposed as a control on rockfall activity by steepening headwalls and increasing fracture density by glacial loading and valley modification
[Miller and Dunne, 1996]. We recognize the importance of these mechanisms in generating rockfall in the Southern Alps; however, if glacial activity is dominant in 28
setting the stage for scree production, scree deposits should be most prevalent along the Main Divide, coincident with the maximum extent of glaciation. Instead, our analyses suggest that the zone of maximum scree extent is offset from the axis of maximum glaciation by ~40 km.
The concentration of active scree within a narrow and distinct altitudinal zone
(Fig. 3-2C) suggests that rockfall activity is controlled by the integrated effects of precipitation and temperature (which vary with elevation and govern vegetation patterns). West of the Main Divide, high pore pressures associated with high rainfall rates have been proposed to enhance bedrock landsliding on steep, heavily vegetated slopes, where [Hovius et al., 1997] documented rapid rates of landslide-driven sediment flux (1.2 × 105 to 1.1 × 106 m3/yr–1). Between 0 and 10 km along our transect a small fraction of the landscape has elevation values above the tree limit
(~1400 m) (Fig. 3-2A), and dense root systems may suppress significant rockfall erosion in favor of deeper bedrock failures. Approaching the Main Divide, mean and maximum elevations increase markedly, such that a large proportion of the landscape is above the tree limit and scree-mantled slopes become more prevalent (Fig. 3-2B).
East of the Main Divide, the proportion of the landscape covered with active scree remains approximately constant (20%–30%), consistent with relatively high maximum elevations and a large proportion of the landscape above the tree limit.
The efficacy of periglacial processes on scree production also depends on altitudinal temperature variations. Recent studies of rock deformation in periglacial 29
environments suggest that the accretion and subsequent preferential freezing of water in large pores may be a dominant factor in rock fracture or displacement (this process is commonly referred to as frost cracking) [Walder and Hallet, 1985]. Theoretical
[Walder and Hallet, 1985], experimental [Hallet et al., 1991], and field-based studies
[Anderson, 1998; Matsuoka, 2001] suggest that frost cracking is efficient within a limited temperature range (−3 to −8 °C) and requires available water. This suggests that where water is readily available, the rate of frost cracking depends on the proportion of time that rock experiences temperatures between -3 and -8°C (Fig 3-3).
As mean annual temperature (MAT) falls below −1 °C , however, a permafrost condition ensues [Anderson, 1998; Ives and Fahey, 1971], limiting the amount of available water and reducing the effectiveness of the frost-cracking process. This
“high and frozen” condition may exist at elevations above 2300 m in the Southern
Alps, consistent with observations of sparse scree and frequent, deep-seated bedrock landslides at high elevations (>2500 m) [McSaveney, 2002]. At MAT above −1 °C (or elevations below 2300 m) frost cracking may be a viable scree production mechanism.
Analysis of annual air temperature fluctuations enables us to quantify the frequency with which rocks of a particular elevation may occupy the frost-cracking temperature window. We averaged and smoothed (using splines) 20+ yr of daily temperature data for five sites in the central Southern Alps, spanning 800 m of elevation (738–1554 m) [NIWA, unpublished data, 2004] (Fig. 3-3). We compared the 30
15 762m 10 C)
o 1554m 5 2300m 0
Temperature ( Temperature -5 Frost Cracking Window
-10 0 60 120 180 240 300 360 Julian Day
Figure 3-3. Annual air temperature variations in the Southern Alps, New Zealand. Temperature data averaged over a period >50 yr at The Hermitage, Mount Cook (762 m, light gray line), was fit with a representative spline (dark line). Using a calculated lapse rate of 0.6 °C/100 m (see text), we reproduced this spline at 1554 m and 2300 m. The recreated spline at 1554 m was compared with 5 yr worth of data collected at that elevation (light gray line) and shows a faithful reproduction of the annual variation at that elevation. At 1554 m elevation, <10 d fit within the frost-cracking window (–3 to –8 °C), compared with >60 d at 2300 m.
difference between smoothed splines at each of these five sites and calculated an average lapse rate of 0.6±0.1 °C/100 m. Using the estimated lapse rate, we calculated the temperature distribution for 2300 m elevation (Fig. 3-3). At 1554 m (the highest
elevation of available temperature data) <10 d fall within the frost-cracking window
during a typical year. In contrast, at 2300 m (coincident with MAT of -1°C), much of
the winter (>60 d) is spent within the frost-cracking zone, suggesting that a small
range of elevations (1600–2300 m) has the potential for sustained frost cracking. 31
Scree deposits form via rockfall erosion of bedrock headwalls, thus the concentration of scree within a limited elevation zone (1200–1600 m a.s.l.) slightly below the frost- cracking window is consistent with frost cracking as the primary control on rockfall erosion.
The pattern of vegetated and active scree represents a temporally integrated signature of rockfall activity (Fig. 3-2B). During the last glacial maximum, the altitudinal zone of efficient frost cracking was considerably lower [Suggate, 1990], such that a significant fraction of terrain in the eastern part of our transect may have undergone frequent rockfall activity. This notion is consistent with our observation that up to 70% of hillslopes in the area reveal scree mantling. By contrast, higher elevation zones near the Main Divide have experienced less time within the altitudinally defined frost-cracking zone (Fig. 3-2B).
In the absence of profound lithologic or seismic controls on rockfall activity, these results highlight the potential for feedbacks between mountain elevation and erosional processes. Whereas previous studies have focused on local relief and glacial equilibrium altitude in assessing topographic controls on denudation [e.g. Brozovic et al., 1997; Schmidt and Montgomery, 1995], here we illustrate the importance of absolute elevation in dictating climatic controls on erosion mechanics.
Quantification of erosion rates in the western portion of the Southern Alps suggests that sediment flux may accommodate tectonic mass influx into the orogen
[Hovius et al., 1997]. Uplift rates of 0.5–1.0 mm/yr for the eastern Southern Alps 32
have been estimated via fission track analyses [Tippett and Kamp, 1993], although these rates remain ambiguous because the effects of lateral advection on fission track ages are difficult to constrain [Batt and Brandon, 2002]. South of our transect, uplift estimates based on the tilting of paleolakeshores reveal similar rates (<1 mm/yr), but their applicability to our study area has not been demonstrated [Adams, 1980].
Despite these uncertainties, available rock uplift estimates in the eastern region of the
Southern Alps approximate estimates of long-term rockfall erosion, suggesting that rockfall erosion may balance tectonic mass flux.
Funding for this work was provided by the National Science Foundation grant
NSF EAR-0309975. We would like to thank Jarg Pettinga, Ben Mackey, and Kerry
Leith for field assistance and discussion. Bill Dietrich, Dave Montgomery and Simon
Broklehurst provided thoughtful reviews that greatly improved the manuscript. Many thanks to Kevin McGill and the National Institute for Water and Atmosphere (NIWA) for access to daily temperature data.
33
CHAPTER IV
ROCKFALL EROSION AND POSTGLACIAL EVOLUTION OF THE
SOUTHERN ALPS, NEW ZEALAND
4.1. Introduction
Uplift and erosion rates vary by order of magnitude across the Southern Alps,
New Zealand [Adams, 1980; Tippett and Kamp, 1993] and are reflected by changes in the dominant style of erosion across the orogen [Whitehouse, 1988]. As the distribution of elevation and precipitation change across the Southern Alps, the relative importance of specific erosional processes change, defining three modern process domains. There is a western domain where high precipitation rates produce a fluvially dissected landscape with high rates of shallow landsliding and ample evidence of larger bedrock landslides (Fig 4-1A)[Griffiths, 1979; Hovius et al., 1997;
Korup, 2006a], an axial domain dominated by glacial processes (Fig 4-1B) [Hicks et al., 1990], and an eastern domain where precipitation and uplift rates are lowest and the landscape is dominated by scree slopes, derived from segregation ice weathering of bedrock (Fig. 4-1C) [Hales and Roering, 2005; Whitehouse, 1988].
A number of studies have suggested that the Southern Alps, New Zealand, are in steady state [Adams, 1980; Batt, 2001; Batt and Braun, 1999; Hales and Roering,
2005; Hovius et al., 1997; Willett and Brandon, 2002; Willett et al., 2001], whereby the rate of tectonic input and resulting uplift of the range is balanced by the amount 34
(A)
(B)
(C)
Figure 4-1. Photographs of differences in the style of erosion across the Southern Alps. (A) An oblique aerial photo of the Arahura valley, West Coast. Note the broad U-shaped valley form that has been deeply dissected. (B) Photograph of Mount Cook showing the Hooker Glacier and the topography typical of the highest parts of the Southern Alps. Here a number of processes are occurring simultaneously including glacial, fluvial and rockfall erosion. (C) Photograph of Purple Hill, Cass Valley in the eastern Southern Alps. The photograph shows the long scree slopes and large debris fans typical of this landscape.
35
of material removed by erosional processes. Authors have argued for a topographic steady state based on similarities between short term uplift and erosion rates
[Griffiths, 1979; Hovius et al., 1997]. Others have proposed a flux steady state, wherein a balance exists between the amount of material entering and exiting the orogen (flux steady state) [Adams, 1980] or an exhumational steady state based on a constant rate of exhumation [Batt, 2001; Batt and Braun, 1999; Tippett and Kamp,
1993]. Studies seeking to decipher the tendency towards steady state conditions in the
Southern Alps have focused on the western side of the orogen because it features highly dissected, landslide-prone terrain, high precipitation rates (>5 m/yr), rapid rates of uplift and erosion (up to 10 mm/yr), and metamorphosed rocks that offer thermochronometric dating opportunities (Fig. 4-2) [Batt, 2001; Griffiths, 1979;
Hovius et al., 1997; Korup, 2005c]. In the eastern Southern Alps, by contrast, erosion and precipitation rates are slower by an order of magnitude, large landslides are infrequent, and the terrain features broad glacial valleys and ubiquitous scree slopes that have been interpreted as being inactive through much of the Holocene
[Whitehouse, 1988; Whitehouse and McSaveney, 1983]. As a result, efforts to estimate the apparent balance (or imbalance) of tectonic and erosional mass fluxes across the orogen have tended to neglect the contribution of erosional mass transfer 36
100 (A) Wellman (1979) Tippett & Kamp (1993)
10
1
0.1 Rock Uplift Rate (mm/yr) Rate Uplift Rock
0.01 0 20406080
100 (B) Hovius (1997) Hales & Roering (2005) Sediment Yields (var. authors) 10 ) r
1
0.1 osion Rate (mm/y osion Rate r E 0.01
0.001 0 20406080 Distance from Alpine Fault (km)
Figure 4-2. Estimates of rock uplift rates (A) and erosion rates (B). Rock uplift is estimated from the offset of geomorphic features and the tilt of lake terraces [Adams, 1980; Wellman, 1979] and from fission track dating methods [Tippett and Kamp, 1993]. Erosion rates have been estimated from mapping of shallow bedrock landslides [Hovius et al., 1997], volume estimates of debris fans [Hales and Roering, 2005] and from sediment yield data [Adams, 1980; Basher et al., 1988; Griffiths, 1979; Griffiths, 1981; Griffiths and McSaveney, 1986; Hicks et al., 1990; Jacobson and Blum, 2003; Thompson and Adams, 1979] 37
originating from the eastern Southern Alps [Griffiths, 1979; Hovius et al., 1997;
Wellman, 1979]. Nonetheless, the scree-dominated, slowly eroding eastern region occupies a much larger fraction of the orogen than the high-erosion western region, such that its contribution to the range-wide flux balance may be significant. Given that glacial-interglacial fluctuations impart a profound topographic signature in the range, their impact on sediment production is likely to be important, although few studies have addressed how such oscillations may affect the apparent balance of erosion and rock uplift rates.
Many recent, process-based geomorphic studies in the Southern Alps provide a rich dataset for exploring the spatial and temporal patterns of erosion across the orogen as modulated by variations in precipitation, rock properties, seismicity, elevation, and rock uplift [Hales and Roering, 2005; Hovius et al., 1997; Korup,
2005b; Korup, 2005c; Korup, 2006a; Korup et al., 2004]. Given the western emphasis of most of these studies, our understanding of controls on erosion in the eastern region remains uncertain. What mechanisms control the transition from glacial to rockfall erosion in the eastern Southern Alps? During interglacial-glacial oscillations, how does the dominant style of erosion change and how does this affect the amount of sediment produced in the orogen? To better quantify regional controls on denudation of the Southern Alps (and in particular the eastern region), this paper expands on previous work by Hales and Roering [2005] by presenting new field data, aerial photograph interpretation, and cosmogenic radionuclide dates that quantify how 38
variations in rock strength, aspect, temperature conditions and glacial-interglacial cycles affect sediment production in the eastern part of the Southern Alps.
4.2. Study Area
The Southern Alps are formed by an oblique continent-continent collision between the Pacific and Australian plates [Walcott, 1998]. Uplift is concentrated along the Alpine Fault, a narrow zone of NNE-SSW oriented oblique thrust faults.
The magnitude of slip on the Alpine Fault varies along strike with highest rates of uplift corresponding with the greatest topographic relief, around Aoraki (Mount
Cook) [Norris and Cooper, 2000]. North of the Alpine Fault, slip on the Alpine Fault is partitioned into a number of smaller strike-slip fault systems (e.g. Hope Fault), before becoming an eastward-dipping subduction zone across the North Island. South of the highest Southern Alps the Alpine Fault moves offshore and becomes a westward-dipping subduction system (Fig. 4-3) [Walcott, 1998].
Rock uplift rates in the Southern Alps have been calculated using the offset of geomorphic features [Adams, 1980; Bull and Cooper, 1986; Norris and Cooper,
2000; Wellman, 1979], extrapolation from thermochronometric dating [Little et al.,
2005; Tippett and Kamp, 1993], and recently, continuous Global Positioning System
(GPS) data [Beavan et al., 2002] (Fig. 4-2). The first attempt to quantify uplift rates for the Southern Alps combined 14C (and other) dating of offset features close to the 39
(A) Elevation (m) (B) 2860 lt au Otira Valley e F o in 168 E lp 0 A 40o S
Main Divide
Cass Valley ns ai Pl y ur rb Christchurch te an C
Figure 4-3. Map of the tectonic setting of New Zealand. (A) Shows the major plate boundary features of New Zealand. Westward-dipping subduction of the Hikurangi Plateau beneath the North Island transitions through a network of faults to the oblique Alpine Fault that defines the western edge of the Southern Alps. South of the Southern Alps the fault transitions to eastward-dipping subduction beneath Fiordland.
Alpine Fault with tilt of lake beds in the southeastern part of the range to produce an uplift map for the whole orogen [Wellman, 1979]; revised by [Adams, 1980] (Fig 4-
2A). The limited geographical extent and poor age control on the lake terraces resulted in a poorly constrained uplift estimate. Dating of terraces on the hanging wall of the Alpine Fault provided uplift rates of ~8mm/yr for 5 sites along the fault [Bull and Cooper, 1986] and 5mm/yr in the Okuru River, South Westland [Cooper and
Bishop, 1979]. The offset of other geomorphic features have been used to calculate
Alpine Fault dip slip rates that vary from >12 mm/yr close to Mount Cook to ~0 mm/yr in the south [Norris and Cooper, 2000]. Exhumation rates of 6 to 8mm/yr have been determined close to the Alpine Fault based on zircon and apatite fission 40
tracks [Batt et al., 2000; Kamp et al., 1989; Tippett and Kamp, 1993], and K/Ar and
40Ar/39Ar dating of muscovites, biotites, potassium feldspars, and hornblendes
[Adams, 1981; Batt et al., 2000; Chamberlain et al., 1995; Little et al., 2005]. These exhumation rates are difficult to relate directly to rock uplift rates due to the effects of erosion and lateral advection. However, analysis of along strike variations in exhumation rate may relate directly to changes in rock uplift rate of ~6 mm/yr
(because of climatic (and erosion rate) consistency across the Southern Alps) [Little et al., 2005]. The lack of reset thermochronometers further than 20km east of the
Alpine Fault make it difficult to extrapolate these exhumation rates across the rest of the mountain range, but some authors have suggested that exhumation rates drop exponentially to ~1mm/yr at the eastern rangefront [Tippett and Kamp, 1993; Tippett and Kamp, 1995]. Recent continuous GPS [Beavan et al., 2002] recordings across the
Southern Alps also support a ~8mm/yr uplift rate close to the Alpine Fault.
The Southern Alps form perpendicular to the major pattern of westerly airflow formed between 40oS and 60oS [Henderson and Thompson, 1999]. This results in a strong orographic precipitation gradient across the Southern Alps, with up to 15 m/yr of precipitation falling close to the midpoint of the western side and decaying exponentially to <1 m/yr in the east [Griffiths and McSaveney, 1983a]. The Southern
Alps have experienced a number of glaciations related to either cooling of global climate [Anderson and Mackintosh, 2006; Suggate, 1990] or strengthening of the westerly airflow across the country [Rother and Shulmeister, 2005]. During the last 41
glacial maximum (LGM) glaciers reached the range front on both sides of the orogen, effectively transporting sediment stored in hillslopes to the Canterbury Plains and sedimentary basins offshore of the West Coast [Gage and Suggate, 1958; Suggate,
1990]. Large depressions in equilibrium line altitude related to at least four glacial advances have been recognized in most major eastward-draining river systems, but the relative age and synchroneity of these glaciations is the source of debate [Almond et al., 2001; Anderson and Mackintosh, 2006; Gage, 1958; Porter, 1975; Shulmeister et al., 2004; Suggate, 1990]. The Otiran glaciation is the youngest, regionally continuous glacial advance and has been correlated to the LGM, with the greatest extent of glaciers being ~18 ka, and with rapid deglaciation beginning at ~14 ka
[Suggate, 1990].
The Southern Alps are almost exclusively composed of interbedded sandstones and mudstones of the Torlesse supergroup. These rocks are turbidites formed in a Jurassic-Cretaceous accretionary prism complex [Wandres et al., 2004].
After deposition, two major orogenies (Rakaia and Kaikoura) uplifted and deformed the rocks, producing complex sets of tectonic joints with no obvious preferred orientations. Tectonic activity has also meant that there is no consistent pattern in bedding orientation or width exists across the range [MacKinnon, 1983]. The only major, consistent change in rock type exists as the amount of regional metamorphism increases from east to west across the orogen, a function of differences in rock uplift rate and exhumation close to the Alpine Fault [Batt and Brandon, 2002]. Within 20 42
km of the Alpine Fault there is a dramatic change from mild pumpellyite-prehnite facies metamorphism to garnet-greenshist and amphibolite facies [Grapes, 1995].
4.2.1. Western Southern Alps
The western Southern Alps have received considerably more study than the rest of the Southern Alps, including sediment yield studies [Adams, 1981; Basher et al., 1988; Griffiths, 1979; Griffiths, 1981; Griffiths and McSaveney, 1983b; Griffiths and McSaveney, 1986; Thompson and Adams, 1979], a bedrock landslide inventory
[Benn, 2005; Korup, 2004a; Korup, 2004b; Korup, 2005a; Korup, 2005b; Korup,
2005c; Korup, 2006a; Korup, 2006b; Korup and Crozier, 2002; Korup et al., 2004;
Korup et al., 2005], and a detailed study of shallow landslides [Hovius et al., 1997] that constrain the spatial and temporal extent of erosion in the range. Initially, sediment yield studies highlighted the very high rates of erosion occurring in the western Southern Alps [Griffiths, 1979; Griffiths, 1981]. These studies showed a power law relationship between precipitation and erosion rate and they calculated a spatially averaged 10 mm/yr erosion rate for westward draining catchments [Basher et al., 1988; Griffiths and McSaveney, 1986], which correlated, within error, with uplift rates of 12 mm/yr [Wellman, 1979] (Fig 4-2).
By mapping the distribution and reflectivity of shallow landslides from two sets of aerial photographs (one flown in 1964/65, the other 1985/86), Hovius et al.
[1997] was able to make a direct comparison between estimated sediment yield 43
produced by shallow landslides and those derived from river-based measurements.
The strong correspondence between these estimates showed that shallow landslides may be responsible for most of the short-term sediment flux in the western Southern
Alps (Fig. 4-2). Recent mapping of landslide deposits and landslide dams has expanded on the work of previous authors [Whitehouse, 1983; Whitehouse and
Griffiths, 1983; and references therein] to show that bedrock landslides affect a large portion of the western Southern Alps [Korup, 2005b; Korup, 2005c]. Large landslides are persistent features in the landscape, but the lack of datable material and their inaccessibility have resulted in poorly constrained estimates of landslide frequency and sediment production [Korup, 2004b; Korup, 2005c; Korup et al., 2004]. By making an assumption about the area-volume ratio of large landslides, Korup et al.
[2004] suggested that short-term post landslide sediment yields could exceed 70,000 t km-1 a-1 (the equivalent of a 33 mm/yr erosion rate), however, the rate of delivery of landslide material to channel systems is unknown.
A suite of mechanisms control high short-term erosion rates in this area including: base level lowering caused by differential uplift across the Alpine Fault
(which causes incision of rivers and destabilization of the surrounding slopes [Hovius et al., 1997; Koons, 1994; Willett, 1999]), pore pressure induced landsliding and rapid sediment removal by fluvial processes [Basher et al., 1988; Griffiths and McSaveney,
1983b; Hovius et al., 1997], and earthquake-induced landsliding and sediment delivery [Korup, 2005b; Korup, 2005c; Korup et al., 2004; Whitehouse, 1983; 44
Whitehouse and Griffiths, 1983]. These erosion datasets for the western Southern
Alps rely on short-term (yearly-decadal) observations of sediment yield and historic landslide mapping, which is problematic given the lack of an historial Alpine Fault earthquake. Without historic seismicity, authors have extrapolated landslide erosion rates based on earthquakes that occurred significantly north or east of the Alpine
Fault [e.g. 1929 Murchison Earthquake; Adams, 1980]. While it is clear that earthquake-induced landsliding is an important process close to the Alpine Fault, debate still exists as to the amount of sediment produced by this process relative to climate-driven slope instability
4.2.2. Axial Southern Alps
Coming from the west, elevations increase significantly close to the Main
Divide, coincident with a major change in erosional processes. The legacy of glaciation is more prevalent, with more obvious U-shaped valleys and a sparse drainage network (Fig. 4-1B). While this alpine terrain is mostly unglaciated, some major glaciers still exist (e.g. Tasman, Hooker). A greater proportion of the landscape can be found above the treeline elevation (~1400 m) [Wardle, 1991]. This treeline elevation is associated with a precipitous drop in the number of shallow landslides
[Hovius et al., 1997] and an increase in the presence of rockfall-derived scree slopes
[Hales and Roering, 2005]. 45
Very few studies have attempted to understand erosion rates in the high
Southern Alps, due largely to its inaccessibility (only 3 major roads cross this region).
Hicks et al. [1990] mapped the input of sediment into the proglacial lake of a retreating Ivory Glacier where rockfall erosion produced a sediment yield of 14,000 t km-1 a-1, which is equivalent to a 6 mm/yr of spatially averaged erosion [using the sediment densities of Griffiths, 1981]. Rockfall erosion rates of 0.3 and 0.6 mm/yr averaged over the last 10ka were estimated using the volume of scree slopes deposited on the sideslopes of the previously glaciated Otira Valley [Hales and
Roering, 2005].
4.2.3. Eastern Southern Alps
The eastern Southern Alps [as defined by Adams, 1980; Whitehouse, 1988] is the most expansive geomorphic domain, yet has received the least study. Depositional landforms are common with sediment stored on the hillslopes as scree slopes and in valleys as large alluvial terraces, alluvial fans, and flood plains (Fig 4-1C).
Large valley systems in the eastern Southern Alps are the legacy of past glaciations which extended to the eastern rangefront [Gage, 1958; Gage and Suggate,
1958; Porter, 1975; Speight, 1916; Suggate, 1990]. The main valleys are wider in the eastern Southern Alps than in the west, which has been attributed to changes in rock strength as a function of metamorphic grade [Augustinus, 1992]. In smaller catchments, the history of glaciation varies systematically from significant glaciation, 46
formed under > 400 ka of glacial action close to the Main Divide to no glacial evidence at the eastern rangefront [Brocklehurst and Whipple, 2004; Brook et al.,
2006; Kirkbride and Matthews, 1997]. All valleys in the range have been highly modified since the last glacial maximum by rockfall erosion and the generation of widespread scree slopes. Moving southeast from the Main Divide, one observes a shift from U-shaped glacial valleys, steep ridgelines, and concave river longitudinal profiles to fluvial valleys with rounded divides, V-shaped valleys, and less concave longitudinal profiles [Brocklehurst and Whipple, 2004; Kirkbride and Matthews,
1997]. Glaciation is characterized by a right-skewed elevation distribution, whereas fluvial processes generate roughly normal elevation distributions with a slight left- skew [Brocklehurst and Whipple, 2004]. Brook et al. [2006] quantified the amount of time required to develop U-shaped valley profiles along the Ben Ohau range, arguing that glacial valleys close to the Main Divide represent at least 400 ka of glacial erosion.
Studies of mass movements in the eastern Southern Alps have concentrated on two extremes of mass movement; very large (>106 m3) rock avalanche deposits
[Burrows, 1975; Orwin, 1998; Smith et al., 2006; Whitehouse, 1981; Whitehouse,
1983; Whitehouse and Griffiths, 1983] and small (<104 m3) rockfall events that create scree-mantled slopes [Hales et al., 2005; Harris, 1975; Whitehouse and McSaveney,
1983]. Analysis of large landslides focus on the slide mechanics of specific events
[Burrows, 1975; Orwin, 1998; Smith et al., 2006] or regional mapping analysis 47
[Whitehouse, 1983; Whitehouse and Griffiths, 1983]. Large landslides are well preserved in this landscape due to the width of the flood plains of the braided river systems, which limits the amount of post-landslide removal of material [Orwin,
1998]. Aerial photograph-based mapping of landslides identified 40 large landslides
(> 106 m3), east of the Main Divide between the Waimakariri and Tasman Rivers
[Whitehouse, 1983]. Volume estimates of these 40 landslides were combined with radiocarbon and weathering rind ages to estimate the frequency of these large events and a spatially averaged erosion rate of ~0.05 mm/yr [Whitehouse, 1983], an order of magnitude lower than other erosion rate estimates for this area [Griffiths, 1981; Hales and Roering, 2005]. Lower precipitation rates on the eastern side of the Southern
Alps have led authors to suggest that many of the very large landslide events are earthquake driven. Two approaches have been used to quantify the mechanisms and rates of rockfall erosion in the Southern Alps, the first used detailed study and dating of individual scree slopes [Whitehouse and McSaveney, 1983], while another took a regional mapping approach [Hales and Roering, 2005]. Whitehouse and McSaveney
[1983] dated boulders on three scree slopes in the eastern Southern Alps using the weathering rind technique and suggested that the slopes accumulated episodically in lobes. Their results showed that scree slopes are modified during and after deposition by debris flow and snow avalanching processes. Hales and Roering [2005] mapped
48
(A)
0 20 40 60 80
a) 80 e (B) r A e 60 op l s l l i 40 f H i
% (
20 e e r c
S 0 0 20 40 60 80 Distance from the Alpine Fault (km)
Figure 4-4. Distribution of scree slopes across the Southern Alps, New Zealand. (A) shows the mapped distribution of active (orange) and vegetated (red) scree slopes. Note the large amounts of scree in the eastern part of the range. (B) Minimum (orange squares) and maximum (red squares) estimates of fraction of hillslopes mantled by scree across our transect. Note rapid increase in scree mantling in eastern part of range. 49
scree slopes from aerial photographs for the transect across the Southern Alps (Fig. 4-
4). Their mapping showed that the scree slopes were most common in the eastern part of the mountain range (occupying up to 75% of hillslope area) precluding earthquakes, rock strength, and topographic unloading as primary agents for creating rockfall erosion in the Southern Alps (Fig 4-4). Instead they noted that scree slopes tend to occur along a narrow elevation range (~1500 m ± 250 m) that was immediately beneath the predicted zone of active periglacial erosion by segregation ice weathering. Segregation ice weathering is an alternative to the traditional view that frost weathering is caused by the change in density when water freezes to ice
(called freeze/thaw). Segregation ice forms as lenses within rock pore space, and causes rock fracture through physical interactions at rock/ice surface [the segregation ice mechanism is discussed in detail by Hales and Roering, 2006; Hallet et al., 1991;
Walder and Hallet, 1985]. Subsequent numerical modeling of conductive heat flow in rock masses further constrained potential limits in the extent of this ‘frost-cracking’ zone to be 1650 to 2000 m, suggesting that the presence of bedrock headwalls and appropriate climatic conditions are a first order control on the amount and rate of rockfall erosion [Hales and Roering, 2006]. 50
4.2.4. Topographic Characteristics
Topographic variations across the Southern Alps provide a simple, yet powerful measure of variations in the style of erosion across the orogen [Adams,
1980; Whitehouse, 1988]. Slope, hypsometry (distribution of elevation), local relief and drainage density characteristics vary systematically within a 40 x 80 km transect oriented perpendicular to the trace of the Alpine Fault (Fig. 4-2). We sampled the distribution of these characteristics in forty 2 x 40 km swaths oriented with the long axis parallel to the Alpine Fault. Major decreases in the median elevation and slope define the change from the eastern Southern Alps to the axial and western areas.
Similarly changes in drainage density define the boundary between the extensively dissected terrain in the west glacially carved valleys along the Main Divide (Figure 4-
5).
Quantile plots of elevation (Fig. 4-5A) show a steep western limb in the
Southern Alps and a much gentler eastern side. West of the Main Divide, the roughly normal distribution of elevation reflects a fluvially-dissected region with long hillslopes and narrow ridges and valleys. Closer to the Main Divide, the elevation difference between the minimum and 5th percentile increases, reflecting greater fluvial incision into glacial valleys (Fig. 4-1B). Proximal to the Main Divide, glacial valleys have been fluvially incised since glacial retreat 10-12 ka [Adams, 1980; Ivy-
Ochs et al., 1999]. Moving eastward, the transition from a glaciated landscape to one 51
2500 (A)
Max 2000 95% 1500 75%
1000 50% Elevation (m) 25% 500 5%
Min 0 0 20406080 350 (B) 300
250
200
150 Gradient (%) 100
50
0 0 20406080 2000 0.08 (C) ) -1 1500 0.06
1000
0.04 500 Mean Local Relief (m) Mean Drainage Density (m Density Drainage Mean
0 0.02 0 20406080 Distance from Alpine Fault (km)
Figure 4-5. The distribution of (A) elevation, (B) slope, (C) local relief and mean drainage density. (A) Plot of the median, minimum, maximum and 5th, 25th, 75th, and 95th quantiles of elevation as a function of distance from the Alpine Fault. Note the dramatic change in median and 25th percentile at about 35 km from the Alpine Fault representing a change to flat, braided river valleys. (B) Plot of the median, minimum, maximum and 5th, 25th, 75th, and 95th quantiles of slope as a function of distance from the Alpine Fault. A change in the median slope also occurs at 35 km from the Alpine Fault. (C) Plot of the mean (dashed line) and standard deviation (grey area) of local relief (5 km diameter window) and mean drainage density (solid line) as a function of distance from the Alpine Fault. Note the dramatic drop in drainage density at approximately 15km from the Alpine Fault that corresponds to a change from the western to axial domains. 52
dominated by braided rivers and scree slopes is represented by a rapid lowering of the median and 25th percentile representing a greater proportion of the landscape at lower elevations (Fig. 4-5A). Similarly the large difference between the median and maximum elevation shows the increase in hillslope length in the eastern part of the range. Such a hypsometric transition represents the change from a predominantly erosional landscape in the west, where sediment is rapidly removed from the landscape by an efficient fluvial system, to one characterized by sediment storage on valley bottoms and on the hillslopes as scree. The distribution of slope along our transect shows a major change at ~38 km, reflecting the transition to greater storage in the east (Fig. 4-5B). Not only is this transition represented by variation in the median and 25th percentile values but by a change in the maximum slope angle. At
~35 km from the Alpine Fault, the maximum value of ~250% (68o) changes to 175%
(60o), reflecting the rounding of ridges and the lack of significant bedrock outcrop in the eastern part of the range.
We measured the local relief along our transect using a 5-km diameter circular window which represents the spacing of the major valleys along the western region.
Mean local relief varies systematically moving away from the Alpine fault, reaching a maximum of 1250 m, at ~5 km from the fault and staying roughly constant until ~25 km from the fault where the mean value decreases with distance from the fault (Fig.
4-5C). This decrease in mean local relief is coincident with an increase in the standard deviation that likely reflects a change in the width of valleys sampled to the 53
west and east of the Main Divide. West of the Main Divide, valleys have a relatively uniform spacing with an average of 5 km measured ridge to ridge. East of the Main
Divide, the landscape is dominated by broader braided river systems (e.g.
Waimakariri, Rakaia, and Rangitata rivers) occupying relict glacial valleys. These main rivers are separated from each other by block faulted ranges (e.g. Cragieburn
Range, Ben Ohau Range). Thus, the mean local relief reflects a change in the width of valleys created by glacial incision and local uplift along faults.
We estimated the drainage density using a local 5-km local sampling window
(for our purposes drainage density is defined as the length of channels occupying a particular area) (Fig. 4-5C). In our study area, drainage density distinguishes fluvial systems from the undissected, broad valleys characteristic of glacial systems [Tucker and Bras, 1998]. Close to the Alpine Fault, mean drainage density is consistently high
(~0.075 m-1), but drops noticeably at the Main Divide. East of the Main Divide, the mean drainage density is highly variable between 0.04 and 0.06m-1 (Fig. 4-5C), due to the influence of large braided river valleys, which have a low drainage density, but vary in width across the central region. In the eastern and central regions, values are consistently lower than other regions because scree slopes inundate the sides of glacial valleys and reduce channel density. The slight increase close to the eastern rangefront reflects a change to greater fluvial incision at the limits of Holocene glaciation (Fig. 4-5C). 54
4.3. Field and Topographic Analysis of the Eastern Southern Alps
Scree slopes in the Southern Alps are found within a restricted range of elevation consistent with rockfall produced by segregation ice driven weathering of bedrock [Hales and Roering, 2005; Hales and Roering, 2006]. We use field-based and topographic analyses to investigate the statistical strength of the elevation trend in scree slopes and the potential role of local terrain aspect in driving rockfall frequency.
In addition, we use cosmogenic radionuclide dates to constrain the frequency of scree production relative to glacial-interglacial fluctuations. We also investigate the distribution of rock strength (and in particular fracture spacing) across our study site to quantify lithologic controls on rockfall generation.
4.3.1. Statistical Analysis of the Elevation of Scree Slopes
4.3.1.1. Methods
Scree slopes have a consistent mean elevation across the Southern Alps which led to the conclusion that they were primarily derived from ice-driven rockfall erosion. A simple test of this interpretation is to compare the elevation of scree slopes to a random sample of hillslope elevations across the same transect. We calculated the mean and standard deviation of scree slope elevation for each of transect swaths (Fig.
4-3). We then created a random sample of hillslope elevations. Based on field observations, we limit our analysis to locales with slope greater than 15%. 55
4.3.1.2. Results
Mean scree slope elevations are uniform and consistently higher than the mean of the random sample in the eastern part of the range (Fig. 4-6). The standard deviation of the random sample is also consistently larger than for observed scree
2000
1600
1200
Elevation (m) Elevation 800
400
0 020406080 Distance from the Alpine Fault (km)
Figure 4-6. Plot of the mean and standard deviation of scree slope (gray shaded zone) and hillslope (dashed line with error bars) elevations. This plot shows that the consistent elevation of scree slopes is statistically significant from a random distribution. 56
slopes. The elevation trend of scree slopes indicates that the slopes may arise from an altitudinally-driven weathering process, such as segregation ice weathering.
4.3.2. The Effects of Aspect
4.3.2.1. Methods
Rock temperature controls the rate of growth of ice in rock masses and therefore the intensity of rockfall erosion [Hales and Roering, 2006]. Due to a lack of rock temperature data in the Southern Alps, we used air temperature as a proxy for rock temperature [Hales and Roering, 2005; Hales and Roering, 2006]. However, rock temperature studies have shown that it is strongly influenced by variations in aspect (or solar radiation input) [Anderson, 1998; Coutard and Francou, 1989; Grab et al., 2004; Hall and Andre, 2001; Hall et al., 2005; Halsey et al., 1998; Ishikawa et al., 2004; Matsuoka, 1994; Matsuoka and Sakai, 1999]. To test whether differences in rock temperatures affect the scree slope distribution, we plotted the aspect of scree slopes across the orogen (Fig. 4-7).
For each 2 x 40 km segment of our transect, we calculated the aspect
(downslope direction) of all scree slope cells using a 25 m digital elevation model.
We filtered the results using a Gaussian low pass filter that removed artifactual aspect measurements, the result of using discrete elevation data in our calculation of slope
57
(A) 10 slopes, N=312 (B) 57 slopes, N=11335 10 km 20 km 0 0
270 90 270 90
180 180
(C) 54 slopes, N= 16911 (D) 19 slopes, N=22169 30 km 40 km 0 0
270 90 270 90
180 180 (E) 55 slopes, N= 33807 (F) 30 slopes, N= 7828 50 km 60 km 0 0
270 90 270 90
180 180
Figure 4-7. Rose diagrams showing the aspects of scree slopes across the Southern Alps. Each rose diagram represents data collected from mapped scree slopes found within a 2x40km rectangle oriented parallel to the Alpine Fault at (a) 10km, (b) 20km, (c) 30km, (d) 40km, (e) 50km, and (f) 60km. Aspect measured from a 25m digital elevation model and the data has been smoothed using a random uniform distribution to remove calculation artifacts from measuring aspect using integer data.
58
10 km 20 km (A) (B) 0 0
270 90 270 90
180 180
(D) (C) 30 km 40 km 0 0
270 90 270 90
180 180
(E) 50 km 60 km (F) 0 0
270 90 270 90
180 180
Figure 4-8. Rose diagrams showing the aspects of hillslopes across the Southern Alps. Each rose diagram represents data collected from a 25m digital elevation model of hillslopes, representing averaged data from (a) 10km, (b) 20km, (c) 30km, (d) 40km, (e) 50km, and (f) 60km distance from the Alpine Fault.
59
values. We then compared the aspect distribution of scree-mantled slopes with the aspect distribution of all hillslopes. To determine the aspect effects of changes in elevation we also made rose diagrams of the orientations of scree slopes as a function of elevation (Fig. 4-9).
4.3.2.2. Results
Aspect varies widely across the study area with no consistency in the mean orientation with distance from the Alpine Fault. At 10km and 60km from the Alpine
Fault, scree slopes are strongly oriented towards the NE and SE respectively (Fig. 4-
7A & 4-7F). There are relatively few scree slopes at these distances from the Alpine
Fault (10km and 60km from the Alpine Fault) and the data is thus biased towards the orientation of the largest scree slopes. The aspects of scree slopes at locations are consistent with aspects of hillslopes in this landscape (Figs 4-7, 4-8). While at some sites a preferred orientation occurs, it typically relates to the dominant topographic oriention of major ridge/valley sequences. For example, at 20km, there appears to be a NW-SE orientation to scree slopes (Fig 4-7B), reflecting a NE-SW orientation of ridgelines (Fig 4-8B). Most importantly, none of these plots suggest that scree production is biased with respect to slope orientation.
Because of differences in the amount of sunlight received by north and south facing slopes over the course of a year we compared the aspect distribution of scree slopes as a function of elevation (Fig. 4-9). The distribution of scree slope aspects 60
0 0 (A) 1000-1100 m (B) 1200-1300 m
315 45 315 45
270 90 270 90
225 135 225 135
180 180
0 0 (C) 1400-1500 m (D) 1600-1700 m
315 45 315 45
270 90 270 90
225 135 225 135
180 180 (E) 1800-1900 m 0
315 45
270 90
0 200 400 600 8001000
225 135
180
Figure 4-9. Plots of aspect as a function of elevation. The aspects of all scree slopes were measured at 100 m intervals measures starting at (A) 1000 m, (B) 1200 m, (C) 1400 m, (D) 1600 m, and (E) 1800 m. 61
does not vary at all with elevation, suggesting that scree is being produced equally from all aspects.
4.3.3. Rock Fracture Spacing
4.3.3.1. Methods
Previous studies have used geotechnical measures of rock strength to characterize rocks in the Southern Alps, Rock Quality Designation (RQD) [Deere and
Deere, 1988] and Rock Mass Strength (RMS) [Moon, 1984; Selby, 1980]. RQD is a common engineering rock strength technique that measures the percentage of a
200cm rock section with segments greater than 10cm [Deere and Deere, 1988]. It is a direct measure of the contribution of fractures to rock strength. The RMS method attempts to account for many rock properties including intact rock strength, amount of chemical weathering of the joint surface, joint orientation relative to the orientation of the slope, joint width, joint continuity, and amount of groundwater outflow [Moon,
1984; Selby, 1982]. After combining these measures, a RMS score between 0 (weak) and 100 (very strong) is assigned to the rock mass.
We conducted scanline measurements of fracture spacing of bedrock headwalls and measurements of the dimensions of rockfall blocks deposited at the top of a scree slope (Fig. 4-10). We measured the long, intermediate and short axes for a random sample of scree blocks found on the upper part of a scree slope, and pooled data for each axis. The data were collected at one headwall/scree site in the 62
Craigeburn Range (43.09341oS, 171.75815oE), and include measurements of both sandstone and argillite beds.
4.3.3.2. Results
The fracture spacing of bedrock headwalls and block size data from scree slopes have mean values of 7.9cm and 4.7cm respectively and are right-skewed (with a skewness of 1.3 and 1.9) (Fig. 4-10). Differences in the mean value and skewness of the distributions is probably related to rock breakage due to impact and the rolling of large blocks down the scree slope [Statham, 1976]. RMS values of Torlesse terrane sandstones and argillites vary between 60 and 95 {Moderate to Very Strong; Moon,
1984], primarily based on the spacing and orientation of fractures as intact rock strength, water flow rate and joint weathering had relatively constant values
[Augustinus, 1995b]. Similarly, RQD measurements show a wide variability from 5% to 79% [Hales and Roering, 2005] (very poor to good rock quality [Deere and Deere,
1988]). The high variability in both the geotechnical measures reflects the bimodal nature of the Torlesse terrane, with highly fractured argillite layers of low strength interbedded with more coherent sandstone layers of greater strength. The range of rock strength and lack of a spatial pattern in the rock strength data suggests that variations in the amount and rate of scree production cannot be attributed simply to a variation in rock type.
63
80 (A) N=144 s t
n 60 e m e r asu e
m 40 of of r e b m 20 Nu
0 0 6 12 18 24 30 Block axis length (all axes)
40 N=107 (B)
30
20
10 Number of measurements Number
0 0 6 12 18 24 30 Fracture spacing (cm)
Figure 4-10. Block size and joint spacing data for the Cass area, Southern Alps. (a) Histogram of the length of all block axes measured at the top of a scree slope above Lake Pearson, Southern Alps (43.09348oS, 171.75810oE). (b) Histogram of fracture spacing measured on the bedrock surface at the location above. 64
4.3.4. Headwall Erosion Rates
4.3.4.1. Methods
We attempted to constrain the frequency of rockfall within the areas of predicted maximum frost cracking rate by measuring the surface age of a bedrock headwall. For areas of high rockfall activity, bedrock headwall surface ages will be uniformly young, whereas area of low rockfall activity will have older surface ages.
We measured the surface age of bedrock headwall segments in the Cragieburn
Range (43.09341oS, 171.75815oE) using a weathering rind method [McSaveney,
1992] as well as in situ 10Be and 26Al. The weathering rind technique has been well calibrated for Torlesse sandstones in the central part of the Southern Alps [Chinn,
1981; Whitehouse et al., 1986]. Using 10Be and 26Al, we calculated the age of five of the oldest intact, in situ outcrops along ridgelines in the Cragieburn Range. Because we were interested in determining the oldest rock headwall ages we biased our exposure age dates by choosing the most highly altered and lichen-colonized surfaces we could find. A preliminary field age was determined using the thickness of weathering rind development, surface color and amount of lichen cover [Bull and
Brandon, 1998; Harris, 1975; Whitehouse and McSaveney, 1983]. 65
4.3.4.2. Results
10Be ages for samples of bedrock headwall age range between 2500 and 0 years. Weathering rind ages have a similar, but slightly narrower age range of 1000 to
250 years (Fig. 4-11). These measured bedrock ages represent maximum headwall ages for the area and are considerably younger than the age of deglaciation for the area [~12 ka; Gage, 1958; Suggate, 1990]. As a result, these slopes are unlikely to only reflect hillslope response to deglaciation. Instead, such young ages suggest that rock headwalls are still actively eroding, consistent with findings from weathering rind dating of scree slope blocks from a scree slope in the eastern Southern Alps
[Whitehouse and McSaveney, 1983]. Weathering rind growth in this system is not affected significantly by proximity to the surface, so the similarity of bedrock and scree ages suggests that headwalls are producing significant amounts of sediment that does not reside on the slopes for long before being covered by freshly produced material. 66
2500 (a)
2000 s) r 1500
1000 Be Age (yea 10
500
0 0 500 1000 1500 Weathering Rind Age (years)
160 (b)
120
80
No. of scree blocks scree No. of 40
0 0 1000 2000 3000 4000 Age (years)
Figure 4-11. Plot of bedrock headwall and scree slope ages, Southern Alps .(a) Plot of the bedrock headwall ages in the Cragieburn Range (43.09341oS, 171.75815oE) calculated using 10Be radionuclide and weathering rind techniques. Note the disparate ages and correlations between both ages. (b) Scree surface ages calculated using the weathering rind technique by [Whitehouse and McSaveney, 1983]. Note the similar age range between bedrock and scree suggesting that slopes are active. 67
4.4. Discussion
The three process domains defined for the Southern Alps represent the dominant erosional processes of the current interglacial, as extensive glaciation during the last glacial maximum produced glaciers that extended to from the range onto the surrounding alluvial plains. The legacy of past glaciations controls the spacing of major river valleys for both westward (e.g. Arahura River, Hokitika River) and eastward (e.g. Rakaia River, Waimakariri River) draining rivers.
Even with these changing conditions related to the expansion and retreat of large ice masses, there is evidence for persistent scree slopes on a thousand year timescale in the eastern Southern Alps. The Southern Alps are home to a unique suite of plants adapted to living on scree slopes [Cockayne, 1958; Edgar and Conner,
1999; Fisher, 1952]. 25 distinct species of plant from 14 plant families grow on scree slopes, rooting deep into the slopes (> 30 cm) where the slopes are stable, finer- grained, and contain more moisture [Cockayne, 1958]. These specialized scree plants are almost exclusively found in areas of Torlesse terrane sandstones and shales
[although some species of grasses (e.g. Poa schistacea) may have adapted to schistose screes; Edgar and Conner, 1999]. These plant species are only found in mountains with elevations greater than 4000’ (1200 m) that receive less than 51” (130 cm) of precipitation [Fisher, 1952] suggesting that scree slopes have been a significant feature in the high Southern Alps for many millennia. These conditions are consistent with parts of the Southern Alps that currently have the highest concentrations of scree 68
slopes and highest rockfall erosion rates [Hales and Roering, 2005] and may provide a first order constraint on the persistence of these features in the Southern Alps.
4.4.1. The Effects of Glacial Cycles on Scree Production
The lack of consistent scree slope orientations at the scale of a 25 m digital elevation model suggests that changes in rock temperature driven by aspect are reflected by relatively minor changes in the amount of scree produced. Mean annual rock temperatures differ from air temperatures by up to 8oC because of the effects of shielding typically caused by trees, snow cover, cloudiness, and aspect which control the amount of sunlight absorbed at the rock surface [Iqbal, 1983]. Many shielding effects are relatively short-lived (seconds to hours), and therefore only affect the shallowest depths in the rock [Anderson, 1998]. Measurements of rock temperature have shown a strong aspect dependence on the minimum, maximum and average rock temperature [Coutard and Francou, 1989; Matsuoka and Sakai, 1999]. Possibly of more importance than rock surface temperature is the amount of snow cover in an area. Snow cover suppresses rock-temperature changes and maintains the rock surface at 0oC [Matsuoka, 2001]. However, most of the headwalls that produce significant rockfall erosion are steep and do not accumulate considerable snow. The lack of a first-order difference in the spatial distribution of scree slopes, particularly a difference between north and south facing slopes suggests that aspect-driven changes in rock temperature do not play a significant role in causing rockfall erosion. It also 69
lends weight to using atmospheric temperature as a proxy for rock surface temperature at the scale of a 25 m digital elevation model.
Post-glacial sediment yield in mountainous environments is thought to decrease exponentially from very high rates immediately after glaciation to relatively low rates currently [Augustinus, 1995a; Ballantyne, 2002]. Evidence for rapid post- glacial sediment accumulation in alpine environments has been interpreted using the volume of slopes preserved in protalus ramparts in Scotland [Ballantyne and
Kirkbride, 1987], the development of early Holocene lake terraces onto Pleistocene debris fans [Arthur, 1975], development of paleosols on debris fans [McArthur,
1987], and the relatively small amounts of observed activity on scree slopes
[Ballantyne and Eckford, 1984; Hinchliffe and Ballantyne, 1999]. There are two competing hypotheses for why this temporal pattern exists. The first suggests that this pattern of sediment yield reflects changes in the rockfall erosion rate through time.
This hypothesis suggests that rockfall erosion rate is primarily due to the fracture and breakup of rock during the topographic unloading of slopes after glaciation
[Augustinus, 1995a; Miller and Dunne, 1996]. If correct, headwall erosion rates should change considerably with time and the surface age of rock headwalls and the material on scree slopes should both have early Holocene ages. Another hypothesis is that the rate of sediment production is a function of the amount of storage on the slope. In this case, sediment is produced from headwalls at a relatively constant rate, but as sediment accumulates on hillslopes, the amount of headwall able to span 70
rockfall decreases [Rapp, 1960; Statham and Francis, 1986]. Our 10Be dating of a bedrock headwall in the Cragieburn Range shows that bedrock headwalls are young
(< 2500 years) and the wide range of ages suggests that rockfall is common (Fig. 4-
10). We suggest therefore that the change in sedimentation rates since the last glacial maximum represent changes in the accommodation space in small drainage basins.
We have shown that the elevation of scree slopes correlates with a zone of predicted maximum frost cracking processes [Hales and Roering, 2006] (Fig. 4-6 and
4-12). This correlation was observed for the distribution of active scree slopes and current temperature conditions found within the range (Fig 4-12A). Considering global temperature changes, one would expect that the distribution of active scree slopes should migrate to lower elevations during glacial periods and higher elevations during interglacials. Hales and Roering [2006] identified a range of elevations in which segregation ice weathering of bedrock is maximized (called the
“frost cracking window”). Their elevation range corresponded to mean annual temperatures between 0 and 2.5oC. Elevations colder than 0oC still crack via this mechanism, but at rates that are an order of magnitude slower than in the frost cracking window due to the effects of permanent ice (a “high and frozen” condition).
At mean annual temperatures between 2.5 and 5oC, segregation ice is present but cracking occurs at a slower rate, and there is no ice growth above 5oC. To explore the implications for range-wide denudation of the Southern Alps, we superimposed a frost cracking map onto the distribution of scree slopes in the Cragieburn Range in 71
(A) Current Temperature Distribution (B) Minus 1o C
N N
1km 1km
(C) Minus 3o C (D) Minus 5o C
N N
1km 1km
Key to temperature ranges < 0o C 0-2.5o C 2.5-5o C >5o C Active scree slopes
Figure 4-12. Shaded relief maps of the Cragieburn Range, Southern Alps. Shown are key temperature ranges for segregation ice weathering for different climatic conditions, (A) shows the current distribution of mean annual temperature, with the other plots showing cooling of mean annual temperatures by (B) -1oC, (C) -3oC, and (D) -5oC. Elevations within the peak zone of frost cracking are designated with red and orange colors. ‘High and frozen’ terrain is white, while unfrozen land is blue. Superimposed on these plots are the distribution of active (cross-hatched) and vegetated (white line) scree slopes. 72
the eastern Southern Alps (Fig. 4-12A). The plot shows a strong correlation between the elevation of active scree slopes and the frost cracking window (in red, fig. 4-12A).
We also show three alternative temperature distributions related to decreases in temperature across the range of -1oC, -3oC, and -5oC (Fig. 4-12B-D), reflecting the range of possible climatic conditions during the last glacial maximum. In each case, as temperatures decrease, the zone of maximum frost cracking moves down in elevation. With a decrease in temperature of -3oC, the elevation of the frost cracking window has dropped by 500 m and much of the upper part of the range is high and frozen (and thus not prone to rockfall via frost cracking) (Fig. 4-12C). Given this amount of temperature lowering (-3oC), the frost cracking window is coincident with the maximum estimated extent of vegetated scree slopes in the Southern Alps (white lines, fig. 4-12). Vegetated scree slopes have been inactive for a long period of time and have developed significant vegetation [Hales and Roering, 2005]. We interpret these slopes as possibly representing scree that was produced during the last glaciation.
4.4.2. Landscape Development on Glacial-Interglacial Timescales
Viewed in the light of changing global temperatures, the erosional development of the Southern Alps varies with distance from the Alpine Fault. In the western Southern Alps, the glacial-interglacial cycle is reflected by periods of glacial incision and downcutting and periods of fluvial incision and landsliding. In the east, 73
interglacial periods are marked by massive sediment production caused by rockfall erosion in the Southern Alps [Hales and Roering, 2005], while little deposition occurs in the Canterbury Plains and offshore basins [Carter, 2005; Leckie, 2001]. Glacial periods are marked by massive accumulations of sediment on the Canterbury Plains, which is typically assumed to be produced primarily by glacial processes, such as plucking and abrasion [Leckie, 2001]. Instead, we suggest that much of the sediment that is deposited on the Canterbury Plains during glacial advance is reworked scree slope and alluvial fan material created during interglacial times. Thus the increase in sedimentation rate reflects the efficiency at which glaciers transport sediment rather than how efficiently they erode bedrock. This suggests that sedimentation in the onshore and offshore basins surrounding the Southern Alps reflects the spatial integration of differing erosion rates across the range and the temporal integration of storage and evacuation cycles that occur at a glacial/interglacial timescale.
4.5. Conclusions
The rate and style of erosion varies across the Southern Alps following variations in precipitation, uplift, and elevation. These variations produce a landscape that can be divided into three different regions, a western region of fluvial dissection and landslides, an axial region of glacial erosion, and an eastern region of braided rivers and scree slopes. While the western region has been well studied, the areally extensive east has received very little work. Hillslopes in the eastern Southern Alps 74
are dominated by scree slopes, deposited by rockfall erosion. We have shown that changes in rock strength cannot produce the necessary pattern of rockfall. Also small- scale variations in rock temperature caused by aspect differences do not appear to affect the distribution of scree slopes. New cosmogenic radionuclide ages show that bedrock headwalls found within the predicted zone of maximum frost cracking erosion are young and actively producing rockfalls. Taken together, these data suggest that rockfall erosion is an active process in the Southern Alps. Furthermore the location of the primary zone of frost cracking is dependent on mean annual temperature such that climatic oscillations should impart significant impacts on orogenic evolution. These data show that to fully understand the history of sedimentation in basins surrounding the Southern Alps requires recognition of the integrated effects of relatively slow rates of erosion acting across large areas of the
Southern Alps. 75
CHAPTER V
CLIMATIC CONTROLS ON FROST CRACKING AND IMPLICATIONS FOR
THE EVOLUTION OF BEDROCK LANDSCAPES
This chapter is in press at the Journal of Geophysical Research- Earth
Surface. I have coauthored this paper with Josh Roering, who provided
advisorial, editorial and financial support for this project.
5.1. Introduction
Sediment production in mountainous regions reflects a complicated feedback between uplift, caused by rock uplift and isostatic rebound, and erosion, which includes the effects of glacial, periglacial, fluvial, and hillslope erosion. Uplift rates change over million-year timescales due predominantly to changes in plate motion and are manifest by broad changes in the sedimentary environment [DeCelles and
DeCelles, 2001]. In contrast, climate change acts over a much shorter timescale and causes rapid pulses of sediment delivery to a basin, reflected by changes in facies, sedimentation rate, and terrace formation [Pan et al., 2003; Zhang et al., 2001]. In high mountains, these changes in sediment delivery have been directly related to the extent of glaciers, which are thought to be the most efficient erosional and transport 76
mechanism in mountains [Brozovic et al., 1997; Hallet et al., 1996; Montgomery,
2002]. Recently, others have suggested that periglacial processes play an important role in erosion, such that glaciers and rivers serve to transport material out of the system [Pan et al., 2003; Zhang et al., 2001]. Quantification of the mechanisms and rates of periglacial erosion is of primary importance to understanding this question.
Figure 5-1. Oblique aerial photograph of the Cragieburn Range, Southern Alps showing large scree-mantled slopes that are the result of active frost processes. 77
The important role that periglacial processes play in mountain erosion is shown by the preponderance of sediment deposited as scree slopes on the flanks of glacial valleys (Fig. 5-1). Enhanced periglacial erosion related to climate cooling events has been suggested as the mechanism causing rapid terrace accumulation events and rapid filling of sedimentary basins [Bull and Knuepfer, 1987; Pan et al.,
2003; Zhang et al., 2001]. Zhang et al. [2001] noted that sediment yields in both offshore and internally-drained basins across the globe show an increase in sedimentation rate at approximately 2-4 Ma, coincident with a change from a relatively stable climate to one with large oscillations on Milankovitch timescales.
The authors proposed that periglacial processes fracture rock and prepare it for transport via glaciers. Similarly, Bull and Knuepfer [1987] suggested that terrace aggradation events are associated with cooler climates in which periglacial processes are most efficient (the study was undertaken in an unglaciated catchment). These studies highlight the need for a process-based understanding of how frost action affects sediment production via the frequency and magnitude of rockfall generation.
Ice-driven mechanical weathering has long been recognized as a causal factor in rockfall erosion. However, previous studies based on a limited number of localized headwalls or scree slopes have focused on the competing importance of ice processes and glacial or topographic unloading [Augustinus, 1995b; Hinchliffe and Ballantyne,
1999], rock properties [Andre, 1997], and earthquakes [Matsuoka and Sakai, 1999].
These studies cover a wide range of climatic and tectonic settings including the 78
European Alps [Sass and Wollny, 2001], the Scottish Highlands [Ballantyne and
Eckford, 1984; Hinchliffe and Ballantyne, 1999], the Japanese Alps [Matsuoka and
Sakai, 1999], Scandinavia [Andre, 1997; Rapp, 1960], the Canadian Rocky
Mountains [Church et al., 1979; Gray, 1973], and Greenland [Frich and Brandt,
1985].
The following examples highlight the diverse character of hypotheses regarding erosion via frost-induced rockfall. The timing of rockfall events and the amount of material accumulated on snow-covered scree slopes was measured by
Matsuoka and Sakai [1999], who suggested that the intensity and duration of freezing and thawing cycles, combined with less frequent, large magnitude earthquake and precipitation events were the primary controls on rockfall of headwalls in the
Japanese Alps. In Svalbard, Norway, a lichenometric study of rock walls of different rock types suggested that lithology was the primary control on rockfall erosion
[Andre, 1997]. Beneath basalt cliffs, according to cosmogenic radionuclide dating of
36Cl in olivine, 80% of rock wall retreat occurred within 6000 years of deglaciation.
This pattern was interpreted as reflecting stress release following deglaciation combined with periglacial processes [Hinchliffe and Ballantyne, 1999].
Measurements of daily crack width changes on a rock wall in the Japanese Alps suggest that ice growth caused an increase in crack width during thawing periods
[Matsuoka, 2001]. A study of the Southern Alps, New Zealand, showed that scree slopes occupy a relatively small elevation range that coincides with areas in which ice 79
growth may be an important weathering mechanism [Hales and Roering, 2005].
Recent work modeling the depth to permafrost in the European Alps, suggested that the thaw of alpine permafrost during a particularly hot summer was responsible for enhanced rockfall erosion rates [Gruber et al., 2004]. Although these studies highlight a diverse dataset of potential mechanisms and have shown that rockfall erosion can occur at geomorphically significant rates [headwall retreat rates of up to
49 mm/yr [Zhu, 1996]], spatial and temporal controls on ice-driven bedrock erosion remains elusive.
Recognition of the important role that ice plays in fracturing and liberating rock from headwalls has led to a number of theoretical and physical experiments that attempt to model ice lens growth [Hallet et al., 1991; Murton et al., 2001; Walder and
Hallet, 1985]. These studies have addressed the mechanics of ice growth, particularly the role of premelted films in lens growth [Dash et al., 2006] and fracturing frequency of rocks under steady temperature gradients [Hallet et al., 1991; Murton et al., 2001]. However, the detail at which these studies have been undertaken makes it difficult to relate the physics of ice growth to rockfall erosion at geomorphic scales.
Here, we provide a theoretical roadmap for predicting the spatial and temporal pattern of rockfall erosion based on changes in climate and rock fracture spacing. 80
5.2. Mechanisms for Rock Fracture by Ice Growth
The physics of ice/rock interactions are a source of debate amongst periglacial geomorphologists, with two competing hypotheses. The first suggests that the ~9% volumetric expansion of ice formed in pore spaces increases tensional stress at crack tips and causes rock to break (hereafter called freeze/thaw). Subsequent thawing liberates the freshly fractured rock, causing rockfall. Freeze/thaw weathering has wide acceptance in the literature [e.g. Church et al., 1979; Coutard and Francou,
1989; Matsuoka and Sakai, 1999], but its applicability to natural systems has been questioned over the past 25 years [Murton, 1996; Walder and Hallet, 1985].
In the 1980s, theoretical [Walder and Hallet, 1985] and experimental [Hallet et al., 1991] studies expanded upon an extensive soil mechanics literature [e.g.
Gilpin, 1980; Taber, 1929] to suggest an alternative mechanism called segregation ice growth. They argued that for freeze/thaw to break rock, at least 91% saturation in a
“closed system” is required to create enough force from ice expansion to break rocks
[Hallet et al., 1991]. In unsaturated or open systems, ice can expand to fill pore space without putting significant pressure on the pore walls. They argued that there are relatively few natural instances in which a completely closed or saturated system exist. Their alternative is a mechanism that can be applied more readily to natural systems involving the migration of water through a frozen fringe towards an ice lens where it accretes (hereafter called segregation ice growth) [Walder and Hallet,
1985]. The segregation ice mechanism proposes that the stress required to fracture 81
rock (or heave soils) is provided by van der Waals and electrostatic forces created at the ice/rock interface [Wettlaufer and Worster, 1995; Wilen and Dash, 1995]. These interfacial forces provide enough stress on fracture walls to cause a small amount of fracture at the crack tip. Once a rock has fractured initially, liquid water is drawn from warmer parts of the rock towards the site of failure along a chemical potential gradient [Wilen and Dash, 1995]. Liquid water exists at sub-freezing temperatures at the interface between rock (or any other solid) and ice due to surface melting and curvature effects [Worster and Wettlaufer, 1999]. These nanometer-scale interfacial water layers [termed premelted films; Wettlaufer and Worster, 1995] vary in thickness as a function of temperature, such that they are thinner at cooler temperatures. These thin, premelted films provide interconnected pathways that allow water to migrate within the frozen zone where it can accrete to an ice lens [Rempel et al., 2004].
The segregation ice mechanism requires that the forces acting at the rock/ice interface (called disjoining forces) are large enough to fracture rocks (or heave soil particles) and that interconnected interfacial water layers allow water to move within the frozen zone. The result is that the efficacy of this mechanism is limited to a small range of temperatures and depths within a rock (or soil column). At temperatures close to 0oC, the disjoining forces are relatively weak although water is readily available because the interfacial layer is thickest. By contrast, at temperatures well below 0oC [experimental evidence suggests <-8oC; Hallet et al., 1991], disjoining 82
forces are strong, but water flow is limited by a thin interfacial layer. In addition, disjoining forces must be strong enough to overcome overburden stresses and cohesive strength of the rock or soil.
In soils, segregation ice growth is largely accepted at the dominant cause of frost heave, due to an extensive history of theoretical and physical experimentation
[Gilpin, 1980; Henry, 2000; Rempel et al., 2004; Taber, 1929]. Ice lenses have grown in porous materials containing a wide range of liquids including water, benzene
[Taber, 1929], helium [Hiroi et al., 1989] and argon [Zhu et al., 2000], some of which contract upon freezing. A similar mechanism may occur in rock masses [Walder and
Hallet, 1985], except that forces occurring at the rock-ice interface must overcome the cohesive strength of the rock mass in order for lens growth to occur. While considerably less theoretical and physical experimentation has occurred on rocks than soils, experiments measuring the location of cracking events [Hallet et al., 1991] and the amount of rock mass heave [Murton et al., 2001] suggest that the segregation ice mechanism may be responsible for a significant amount of fracturing in rock masses.
This is further supported by field experiments measuring the change in crack width with time in the Japanese Alps, which showed that the peak crack width coincided with a peak in water availability in the spring [Matsuoka, 2001].
For geomorphologists, the most important difference between the segregation ice growth and freeze/thaw mechanisms is the temperature conditions over which they operate. Freeze/thaw weathering requires temperature oscillations about 0oC, 83
whereas segregation ice grows at a temperature range that maximizes the availability of water and the strength of the surface forces needed to crack rock. Theoretical and experimental work suggests that the growth of segregation ice lenses depends primarily on absolute temperature (the growth rate is largest at -3 to -8oC), the amount of available water, and a temperature gradient that allows water to be drawn from warmer (>0oC) to cooler regions of the rock [Hallet et al., 1991; Matsuoka,
2001; Murton et al., 2001; Taber, 1929; Walder and Hallet, 1985; Wilen and Dash,
1995; Zhu et al., 2000]. These differences between hypothesized mechanisms are critical for understanding mountain evolution, particularly the elevation at which sediment is likely to be produced and the temperature conditions associated with rock mass failure. If freeze/thaw is the dominant mechanism, maximum frost cracking intensity (and rockfall erosion) should occur at an elevation that experiences the greatest number of oscillations across 0oC. If the segregation ice growth mechanism dominates, the maximum frost cracking intensity should occur at an elevation that spends the most amount of time within a temperature range of approximately -3 to -
8oC (in the presence of available water).
To explore the role of segregation ice growth on mountain evolution we present a simple one-dimensional heat flow model that describes how the intensity of segregation ice growth varies with depth. Our model extends the work of Anderson
[1998] who used an analytical solution to understand the number of days a rock spends within a temperature controlled window (which he termed the frost cracking 84
window). We have extended his work to account for hydrologic and temperature gradient factors. Our model is based on an analytical solution for heat flow in a solid mass with a sinusoidal, annually varying upper boundary condition [derivation of this solution is described in Carslaw and Jaeger, 1959; application to permafrost systems is described in Gold and Lachenbruch, 1983].
5.3. Model Setup
We modeled shallow bedrock heat flow in an attempt to predict the depth at which segregation ice is likely to form in a rock wall. Conduction is the primary method of heat transfer in rocks at Earth’s surface [Turcotte and Schubert, 2002].
Neither radiative or convective cooling are thought to play major roles in heat transfer within solid rock, but rocks with large interconnected pore spaces filled with fluid may experience a component of convective cooling [Harris and Pederson, 1998]. In the case of highly fractured rock the relative roles of conduction and convection have not been assessed, but we assume that conduction dominates [Anderson, 1998; Gold and Lachenbruch, 1983]. The analytical solution for one-dimensional heat conduction given a sinusoidal surface temperature variation is
π − z αP ⎛ 2πt π ⎞ T(z,t) = MAT + Tae y sin⎜ − z ⎟ , ⎜ P αP ⎟ ⎝ y y ⎠ where T is temperature, t is time, z is depth in the rock, MAT is the mean annual temperature, Ta is the half amplitude of the sinusoidal variation, Py is the period of 85
the sinusoid (our model uses an annual cycle) and α is the thermal diffusivity
(between 1-2 mm2 s-1 for most rocks) [Anderson, 1998; Carslaw and Jaeger, 1959;
Gold and Lachenbruch, 1983]. This model predicts that the daily temperature cycle is attenuated by an order of magnitude at 35 cm and the annual temperature cycle is comparably attenuated to a depth of 700 cm [Gold and Lachenbruch, 1983].
We calculated the location and efficacy of segregation ice growth using three criteria. The first is that the rock temperature at which segregation ice grows is between -3 and -8oC [Hallet et al., 1991]. Second, water (T>0oC) is available to the system, which corresponds to scenarios with melting conditions at the surface, or unfrozen groundwater at depth. In our model, water is available when either boundary has a temperate above 0oC. Finally, when water is available to the system from groundwater (defined as when the lower boundary at 20 m depth is above 0oC), segregation ice grows when the temperature gradient is positive (i.e. water is drawn from warm to cold). Alternatively, when water is available at the upper boundary
(during times of melting), segregation ice grows when the temperature gradient is negative.
In an attempt to directly relate our model to frost cracking processes we use the temperature gradient as a proxy for frost cracking intensity. Ice segregation theory suggests that the growth rate of segregation ice varies with the temperature gradient
[Worster and Wettlaufer, 1999]. A steeper temperature gradient reflects a steeper gradient in chemical potential and therefore more rapid delivery of water to sites of 86
ice accretion. The rate of ice lens growth is primarily dependent on water availability
[Wettlaufer and Worster, 1995], so more efficient delivery of water to the system will result in faster lens growth, greater stress at crack tips, and more intense frost cracking at that location. We suggest therefore that by summing the temperature gradient at every predicted location over the course of an annual cycle we can quantify the rate of segregation ice growth. Because of a strong predicted relationship between the growth rate of ice and frost cracking intensity [Walder and Hallet, 1985], we have used the annual sum of the temperature gradient as a function of depth as a proxy for the intensity of frost cracking.
We solved equation (1) at daily intervals for a year with Ta varying between 4 and 20oC and MAT varying between -5 and 5oC, reflecting the variation in MAT and variability of mountainous and arctic areas around the world [Anderson, 1998]. At depths below the penetration of the maximum annual cycle, we kept the rock mass at a constant temperature consistent with the mean annual temperature. Because the maximum depth of frost penetration was ~14 m, the effects of geothermal heating are negligible [a typical geothermal gradient is ~0.025oC/m, Turcotte and Schubert,
2002] when compared with surface temperature changes.
5.4. Model Assumptions
We attempted to model the complicated physics of segregation ice growth using readily accessible parameters, such that several assumptions are embedded in 87
our calculations. The first is that the segregation ice mechanism is most effective in rock at temperatures between -3 and -8oC. As we have discussed above, the efficacy of segregation ice growth is also highly dependent on the cohesive strength of the rock mass and the overburden pressure at the location of ice growth.
The second important assumption regards the availability of water in our model system. Water is available at the surface in the spring during thawing periods or from a groundwater system that is above 0oC. In the spring, one often observes considerable water along the surface. These observations are supported by direct measurement of crack widening in the spring thaw in the Japanese Alps [Matsuoka,
2001]. The presence of groundwater in steep alpine cliffs is more difficult to justify as few measurements have been made due to the difficulty of instrumenting these areas.
A recent study used a number of geophysical techniques to show that the moisture content of rocks in the European Alps varied considerably, from 50 to 100% at 5cm depth to a nearly uniform 80% at 10 cm depth [Sass, 2005]. The results of this study suggest that groundwater is present in alpine cliffs and thus may be available to promote segregation ice growth, although the rate of water movement is dependent on the permeability of the interconnected pores and the magnitude of the chemical potential gradient.
Our model focuses on the annual temperature cycle and neglects diurnal effects. The diurnal variation is important only in the top few centimeters of the rock mass, as it is attenuated by an order of magnitude at a depth of 35 cm. Although this 88
may be an important component in some areas, the effects of diurnal variations are relatively small compared with those of the annual cycle. The assumption of a sinusoidal annual temperature variation provides the simplest reasonable approximation of a natural system. In detail, the annual cycle approximates a sinusoid during cooling periods (autumn), but natural temperature rise much more slowly during warming periods (spring). Air temperature close to Earth’s surface is primarily dependent on the sun’s energy reradiated from the ground [Linacre and Geerts,
1997], so slow warming of Earth’s surface in the springtime causes this anomaly.
However, the effect of this small discrepancy within our model is relatively minor.
We assume that atmospheric temperature is representative of rock surface temperature. The relationship between rock temperature and atmospheric temperature is affected by rock type, albedo and aspect [Hall et al., 2005]. Rocks that have a low albedo and/or are exposed to long periods of solar radiation typically reach temperatures significantly higher than atmospheric temperature for seconds to minutes [Anderson, 1998]. This can result in a discrepancy between atmospheric
MAT and a MAT measured at the rock surface of up to 8oC [Coutard and Francou,
1989]. However, the surface temperature of fully-shaded rocks is similar to measured atmospheric temperatures [Matsuoka and Sakai, 1999]. Our simple model does not attempt to capture all of the variability in rock temperature, but uses atmospheric temperature as a proxy for the shade temperature of rock. 89
This model seeks to understand potential locations of segregation ice growth and does not consider the saturation, permeability or porosity of the rock. We assume that water can move readily, and is limited only by temperature effects. Segregation ice growth in rocks will, in reality, be limited by differences in rock properties, particularly cohesive strength, fracture spacing, and thermal properties. Of these properties, fracture spacing is possibly the most important control on porosity, permeability and cohesive strength of rocks [Selby, 1993]. These assumptions are necessary for the simplification of our model and suggest that our predictions represent maximum values for the depth and cracking intensity of rock masses.
5.5. Model Results
We determined the depth, timing and intensity of segregation ice growth for a range of MAT and annual temperature variations (Ta) that represent temperature conditions commonly observed in nature. MAT primarily controls the availability of water while Ta controls the depth of segregation ice growth and cracking intensity.
5.5.1. Timing of Segregation Ice Growth
The timing of segregation ice growth reflects the interaction between the mean rock temperature and the temperature of the upper boundary condition (i.e. time 90
MAT=0.001oC, Ta=12oC (A) (B) Temperature (oC) Temperature Gradient (oC/cm) -15 -10 -5 0 5 10 15 0 0.010.020.030.04 0 Winter (t=273) 0 Summer (t=91) t=273
400 t=1 Spring (t=1) 400 Autumn (t=182)
800 800
Depth (cm) 1200 1200
1600 1600 Frost Cracking Window
2000 2000
Figure 5-2. Results for our heat flow model at MAT of 0.001oC, Ta of 12oC. (A) Plot of some representative temperature profiles at time steps 1 (spring), 91 (summer), 182 (autumn), and 273 (winter). Highlighted in solid black (day 1) and dashed lines (day 273) are examples of when our model will predict segregation ice to grow. The intensity of ice growth is primarily a function of temperature gradient and its daily values are shown in (B).
of the year). During the autumn and early winter, the surface temperature gets cooler, creating a positive temperature gradient close to the surface that draws water from the warmer interior of the rock towards the surface (e.g. Fig. 5-2a, time step 273). This suggests that the segregation ice growth is likely to occur only in rock masses that have available groundwater (positive MATs). In the late winter and spring, the surface warms such that the temperature-depth profile has a negative temperature 91
MAT=-5oC, Ta=12oC
Temperature (oC) (A) Temperature Gradient (oC/cm) (B) -20 -15 -10 -5 0 5 10 0 0.004 0.008 0.012 0.016 0.02 0 0 Winter (t=273) Summer (t=91)
400 400 Spring (t=1) Autumn (t=182) t=91
800 800
Depth (cm) 1200 1200
1600 1600
2000 FrostWindow Cracking 2000
Figure 5-3. Results for our heat flow model at MAT of -5oC,Ta of 12oC. (A) Plot of representative temperature profiles at time steps 1, 91, 182, and 273. Highlighted in solid black (day 91) is an example of when our model will predict segregation ice to grow. The intensity of ice growth is primarily a function of temperature gradient and its daily values are shown in (B).
gradient close to the surface that allows water to be drawn from the surface, suggesting segregation ice growth for negative MATs (e.g. Fig. 5-3a, time step 1).
For regions with positive MAT, segregation ice grows in the autumn, winter and spring. In this case the lower boundary condition (the internal part of the rock) is always above 0oC, such that water is available from the groundwater system in the warmer interior of the rock. Thus, the number of days during which segregation ice 92
grows is a function of the number of days with a positive temperature gradient at depths where the temperature is between -3 and -8oC (hereafter called the frost cracking window) (Fig. 5-2). In the autumn and early winter (between days 182 and
273 in Fig. 5-2) the temperature at the surface decreases rapidly, creating large positive temperature gradients close to the surface. These steep temperature gradients cause rapid segregation ice growth during the autumn (Fig. 5-2, time step 273).In the late winter and spring, the surface temperature increases, causing negative temperature gradients at the surface and stopping nearby ice growth. However, there is a lag between temperature at the surface and interior depths, resulting in deep ice growth in the spring (Fig. 5-2, time step 1).
Segregation ice growth is limited for negative MAT areas because water is only available to the system when the surface temperature is above zero (Fig. 5-3a, time step 91). In this case, segregation ice grows in an active layer, where water is available from snow melt and follows a negative temperature gradient. The contribution of snowmelt and the formation of an active layer only occur in the spring and summer months.
5.5.2. The Influence of Mean Annual Temperature
Mean annual temperature affects the overall behavior of this system by controlling the availability of water (Fig. 5-4). Changes in MAT cause both the depth of ice growth and the depth of maximum cracking intensity to deepen. For positive 93
MAT sites, the maximum depth of segregation ice growth and maximum cracking intensity gets deeper with decreasing MAT. For example, at a MAT of 5oC the
Ta=12oC Cracking Intensity (oC/cm) (A)Cracking Intensity (oC/cm) (B) 012340 0.2 0.4 0.6 0.8 0 0
o C T=5 MA o C T=4 MA o C MAT=-1 T=3 200 MA o C 200 2 MAT=0 T= MA o C 1 o = C o T 1 C A 00 M 0. MAT=-5 = o T C A M 400 400 MAT=-4 o C MAT=-3 o C MAT=-2 Depth (cm) Depth 600 600 o C o C
800 800
1000 1000 Positive MAT Negative MAT
Figure 5-4. Changes in cracking intensity as a function of MAT for Ta of 12oC. (A) Variations in the cracking intensity and depth of segregation ice growth for positive MATs. (B) As (A) for negative MATs.
maximum depth of ice growth is 120 cm, with the maximum cracking intensity located at the surface, whereas for MAT equal to 0.001oC (representing a minimum positive MAT, at 0oC, water is not available from groundwater) the maximum ice growth depth is 430 cm and the depth of maximum cracking intensity is 150 cm. The 94
difference in the depth of maximum cracking intensity between MAT of 0.001 and
5oC reflects the proportion of the year our model predicts cracking to occur (Fig. 5-5).
At 0.001oC, the surface temperature varies between -12 and 12oC (as compared with
Ta =12oC (A) (B) Number of Days Number of Days 04080120 04080120 0 0
MAT=5oC MAT=4oC 200 MAT=3oC 200 MAT=2oC o MAT=0 C o C o -1 MAT=1 C =
T MAT=-3 C A o M 2 M o M - A 400 MAT=0.001 C 400 A = T = T T - A = 5 o - o C
C 4 M o C
Depth (cm) Depth 600 600
800 800
1000 1000 Positive MAT Negative MAT
Figure 5-5. The number of days spent in the frost cracking window as a function of depth for (A) positive and (B) negative MAT. The data in (A) show the predicted frost cracking window estimated by Anderson [1998] (recalculated from his paper).
-7 and 17oC for a MAT of 5oC) and diffusion of heat from the surface allows deeper rock to enter the frost cracking window. The overall shape of the frost cracking curve
(Fig. 5-4) somewhat resembles profiles of the number of days spent within the frost 95
cracking window (Fig. 5-5). The difference arises from the influence of the temperature gradient, which is steepest close to the surface such that predicted cracking intensity is higher close to the surface than at depth. For higher MATs, which barely penetrate the -3 to -8oC temperature window (e.g. 4 and 5oC) the maximum cracking intensity is close to the surface. When the surface temperature cools below -8oC, the depth of maximum cracking intensity moves away from the rock surface and decreases in peak intensity.
Negative MATs show a much more complicated response to changes in MAT because water is only available for limited periods during the year. The cracking intensity curves show: (1) a broad peak centered between 200 and 300 cm, (2) negligible cracking within 80 cm of the surface, and (3) a long tail that extends deeper than 1000 cm (Fig. 5-4b). Also, the peak cracking intensity is an order of magnitude less than for positive MAT sites. The depth pattern of cracking intensity for MAT just below zero reflects the number of days available for segregation ice growth (c.f. Figs
5-4b & 5-5b). For lower temperatures (between -3 and -5oC) there are two distinct inflections in the cracking intensity distribution (Fig. 5-4b). The upper broad peak corresponds to the depth with the steepest average temperature gradient and reflects the relatively large number of ice growth days. A second inflection point occurs at depths below 500 cm, reflecting a rapid decrease in the number of days at which cracking occurs (due to relatively short periods of time spent with a negative temperature gradient). 96
An important feature of these simulations is the sharp transition in behavior between 0.001oC and 0oC. This transition is a result of the assumption that water is available from the groundwater system at MATs above 0oC. In our model, this transition represents the change from a seasonal ice condition to one of permanent ice
(permafrost). While the abruptness of this change may be difficult to justify in natural systems, the transition between seasonal ice and permafrost is real, the implications of which are discussed below.
5.5.3. The Influence of Annual Temperature Variations
Increasing Ta affects the rate at which the surface temperature passes through the -3 to -8oC temperature window and also the magnitude of changes in temperature at depth. For sites with a positive MAT, an increase in annual temperature variation causes the maximum depth of segregation ice growth and the depth of maximum cracking intensity to deepen (Fig. 5-6a). The change in the maximum depth of segregation ice growth reflects variations in temperature at depth within the rock. For example, a change from an annual temperature variation at the rock surface from 4oC to 20oC represents a change from 0.4oC to 2oC at 7m depth. Changing the amplitude of the temperature variation increases the depth at which the rock crosses into the frost cracking window. Also, an increase in the amplitude of the annual temperature variation results in shallower levels of the rock spending more time below -8oC, where ice growth does not occur. 97
Negative MATs produce a more complicated cracking intensity-depth profile than positive MATs (Fig. 5-6b). Increasing Ta causes a decrease in the minimum depth of segregation ice growth, as well as a flattening and widening of the zone of peak cracking intensity (Fig. 5-6b). Shallowing of the minimum ice growth depth occurs because the first icy day occurs earlier in the year (earlier in spring) when surface temperature gradients are steeper and the -3oC isotherm is shallower. The
(a) MAT = 0.001oC (b) MAT =-5oC Cracking Intensity (oC/cm) Cracking Intensity (oC/cm) 0123 0 0.2 0.4 0.6 0.8 1 0 0
Ta=4oC
200 200
Ta=8oC 400 400 Ta=12oC o o Ta=8 C Ta=16 C o
Depth (cm) Ta=12 C 600 600 Ta=20oC
800 800 Ta=16oC
Ta=20oC 1000 1000 Negative MAT Positive MAT
Figure 5-6. Changes in cracking intensity with depth for different Ta and a MAT of (A) 0.001 and (B) -5oC. 98
flattening and broadening of the peak in maximum cracking intensity reflects increased penetration of the surface temperature perturbation into the rock mass. As a result, more days are spent within the frost cracking window at depth (usually between 400 and 600 cm) but temperature gradients are lower at that depth, leading to a flat profile (c.f. Figs 5-5 & 5-6).
5.5.4. Summary Effects of Ta and MAT
To apply our model to natural systems, we have summarized our depth profiles by using simple measures of the depth, variability, and intensity of frost- cracking conditions. We have chosen four quantities to describe our model results; mean depth of ice growth, maximum depth of ice growth, standard deviation of ice growth depth, and maximum cracking intensity. These quantities best describe the range of variability in depth and intensity of segregation ice growth given the differences in the depth profiles of cracking intensity between positive and negative
MAT. Profiles for positive MAT sites extend from the surface with a sharp peak in cracking intensity at or close to the surface. In contrast, negative MAT sites have a broad peak in cracking intensity at depth within the rock that does not reach the surface. The difference in shape of these distributions suggests that while much of the variability in ice growth depth can be represented by a maximum for positive MATs, the mean and standard deviation are more appropriate for negative MAT sites. 99
For positive MAT sites, maximum ice growth depth decreases systematically with increasing temperature (Fig 5-7a). Maximum ice growth depth is less useful for characterizing segregation ice depth for negative MATs because segregation ice growth can occur at very low rates deep within the rock mass (up to 1200 cm depth)
(e.g. Fig. 5-6), therefore we use the mean ice growth depth and its standard deviation
(Fig. 5-7a & b). Mean ice growth depth for negative MATs is relatively consistent between 250 cm (for warmer MATs) and 450 cm (for cooler MATs) (Fig. 5-7a). The largest mean depths occur when Ta is small, but the mean depth rapidly achieves consistency with increasing temperature variation. Consistency in mean ice growth depth contrasts with large changes in the standard deviation of depth as a function of
Ta and MAT (Fig. 5-7b). The standard deviation increases to reflect the flattening of the distribution with increasing Ta (Fig. 5-6b).
An order of magnitude difference in cracking intensity exists between the positive and negative MAT regions (Fig. 5-7c). Rock masses with positive MATs experience a maximum cracking intensity that increases from 3 to 4oC/cm between
0.001 and 5oC at Ta above 12oC. The finite number of days a rock mass can spend within the -3 to -8oC temperature window regulates the maximum cumulative gradient that can be achieved for a given MAT. Higher MATs experience greater maximum cumulative temperature gradients because they cross into the temperature window during the late autumn and winter, close to the minimum value in the annual temperature curve, when daily temperature changes at the surface are small. A useful 100
600 600
400 400
200 200 Ta = 4oC rost cracking depth (cm) depth cracking rost f Ta = 8oC Ta = 12oC Ta = 16oC
Mean Ta = 20oC
0 0 (cm) depth Maximum cracking frost -5 -4 -3 -2 -1012345 0
300 (B) 300
200 200 Cracking Depth (cm) Depth Cracking 100 100 f St. Dev. o Dev. St. St. Dev. of Cracking Depth (cm) Depth St.Cracking of Dev. 0 0 -5 -4 -3 -2 -1 0 1 2 3 4 5
4 Ta = 6oC (C) Ta = 10oC C/cm) o 3
2
1
MaximumCracking Intensity ( 0 -5 -4 -3 -2 -1 0 1 2 3 4 5 Mean Annual Temperature (oC)
Figure 5-7. Summary plots of the effects of changes in MAT and Ta. (A) Changes in the maximum (for positive MATs) and mean (for negative MATs) cracking depth with MAT and Ta. Positive MATs show a progressive increase in the maximum depth of frost cracking with decreasing MAT and increasing Ta. Negative MATs show an increasing mean cracking depth with decreasing MAT but relative uniformity with variations in Ta. Mean cracking depth was calculated as the mean of the distribution of cracking intensity as a function of depth. (B) Standard deviation of cracking depth as a function of MAT. The plot shows that the width of the distribution increases with increasing TA for both negative and positive MATs. (C) Maximum cracking intensity as a function of MAT and Ta. The peak cracking intensity is an order of magnitude greater for positive MAT. 101
illustration of how maximum cracking intensity is related to days spent in the cracking window is shown for Ta values below 12oC. At low Ta, the pattern of increasing maximum cracking intensity with MAT is reversed because higher MAT areas spend very few days in the temperature window. For example, at a Ta of 6oC and a MAT of 0.01oC, a rock mass spends 122 days in the temperature window, whereas regions with MAT of 2.5oC experience only 48 days. The maximum cracking intensity calculation is less meaningful for negative MATs as the distribution of cracking intensity flattens with higher Ta. Importantly, the maximum cracking intensity is considerably smaller for negative MAT.
5.6. Discussion
Our model results predict the depth of potential segregation ice growth in a rock mass based on mean annual temperature and annual temperature variation. This model can be coupled with relevant geotechnical data such as rock fracture spacing and rock strength measurements (such as Rock Quality Designation [Deere and
Deere, 1988] and Rock Mass Strength [Selby, 1980]) to predict patterns of ice-driven rockfall erosion across mountain ranges. Advantages of this model are its simplicity and that it uses readily available data. 102
5.6.1. Segregation Ice Growth and Rock Fracture
For segregation ice growth to be an important geomorphic agent it must be efficient at fracturing bedrock. Rocks typically contain fractures formed by a number of topographic, depositional and tectonic processes, which commonly provides most of ths primary porosity and permeability of a rock mass [Selby, 1982]. Both porosity and permeability are important parameters in segregation ice growth as pores in rock provide sites for the nucleation and growth of ice crystals, and permeability controls the efficiency with which water can be drawn from warmer parts of the rock. The model results described above show that segregation ice growth is most efficient above 300 cm for positive MATs and above 700 cm for negative mean annual temperatures. These results suggest that rocks with significant porosity and permeability above these depths may be subject to weathering via the segregation ice mechanism.
While it is not clear if frost processes create primary fractures or if they enlarge existing fractures, ice growth is thought to correlate with fracture properties such as permeability, porosity, and fracture spacing [Hallet et al., 1991; Murton et al.,
2001; Walder and Hallet, 1985]. The shear strength and orientation of fractures controls the overall strength of rocks [Deere and Deere, 1988; Hoek, 1983; Selby,
1982]. The shear strength of a fracture is controlled by its roughness, cohesion from intermolecular bonding and the amount of clay minerals present within the joint. 103
Segregation ice growth reduces the shear strength of fractures by creating a tensional stress, which can break intermolecular bonds, dilate the rock mass, and reduce the effectiveness of joint roughness. As such, rock fractures are relevant to mountain erosion via frost processes because they serve as the loci of ice growth as well as provide planes of weakness over which failure can occur. In porous systems, particularly soils, heaving pressures associated with water migration may exceed 20
MPa, ample pressure to cause rock fracture [Walder and Hallet, 1985]. Walder and
Hallet [1985] coupled a model of segregation ice growth with a rock fracture model of stress in “penny-shaped” cracks, and showed that new fractures propagate up to 10-
7 m/s at an ice pressure of 10-12 MPa. Experiments of frost cracking in unfractured sandstones confirmed these theoretical predictions [Hallet et al., 1991]. An important assumption in these models is that ice growth controls the development of new fractures which we view this assumption as an end-member in a natural system.
We suggest fracture spacing may provide a simple, easily measured proxy for determining how susceptible a rock mass is to erosion by segregation ice growth.
Rock fractures that intersect our predicted zone of segregation ice growth are therefore susceptible to frost cracking. We suggest that coupling our model with field measurements of the fracture spacing may provide a first-order understanding of likely sites of rockfall erosion. 104
5.7. Comparison to Previous Studies
In an attempt to provide some initial constraints on our simple model, we compare our modeled results with three different datasets from the literature.
5.7.1. Timing of Rockfall Erosion
A test of the applicability of our model is to compare our estimates of the timing of segregation ice growth with the timing of rockfall erosion. Our model predicts that rockfalls are more likely in the autumn and early winter for locations with a positive MAT and in the spring and summer for locations with a negative
MAT (Figs 5-2 & 5-3). Two rockfall inventories in the Canadian Rockies counted the number of rockfall events in a small alpine drainage basin and showed that rockfall frequency peaked in the spring, with a smaller peak in the autumn [Gardner,
1983; Luckman, 1976]. While no accurate temperature data exists for the sites considered, interpolation of MAT from nearby temperature recordings coupled with a lapse rate of 0.6oC/100m show that both sites have a MAT of less than 0oC. A study of the volume of rockfall debris falling onto a snow covered talus slope in a cirque in the Japanese Alps (MAT of ~-2oC), showed a similar pattern of rockfall frequency, with a large peak in the spring, and a smaller peak in the autumn [Matsuoka and
Sakai, 1999]. These field observations appear to support a peak in rockfall activity during spring thawing periods for rocks found in areas with a MAT < 0oC, however the also show a peak in activity in the autumn, which is not consistent with our 105
model. A study of the amount of rockfall debris collected over the between December
1983 and April 1984 in the Niagara Escarpment, Canada (MAT of 6oC), showed a peak in rockfall activity in the February and March [Fahey and Lefebure, 1988]. The peak in these data is also not entirely consistent with predictions made in our model
(we would predict rockfall earlier in the winter. The discrepancies between our model predictions and the timing of rockfall may reflect the influence of diurnal temperature fluctuations, which will particularly affect densely fractured rock masses.
5.7.2. Southern Alps, New Zealand
Hales and Roering [2005] attempted to understand the rate and distribution of rockfall erosion in the Southern Alps by mapping rockfall deposits (scree slopes) in a transect across the orogen. Scree slopes in the Southern Alps are well preserved by the geometry of the glacial valleys in which they are deposited and thus accurately preserve a record of rockfall erosion. The authors showed that the fractional area of scree-covered slopes peaked in the eastern part of the range, which precluded earthquakes, topographic unloading, and rock type from being the dominant rockfall generation mechanism. Instead, scree slopes were found within a narrow elevation range with a mean of ~1500 m and a standard deviation of ~250 m, coincident with elevations immediately below the frost cracking window.
We used their dataset to test our model by comparing the predicted elevations of maximum intensity segregation ice growth with the elevations of observed rockfall 106
erosion. We calculated the annual temperature variation and MAT as a function of elevation and using 7 climate stations, found between 42.5 and 44oS latitude and covering an elevation range of 762 to 1554 m over 20 to 99 years [NIWA, unpublished data, 2004]. All climate stations showed a narrow annual temperature variation of 12oC (Ta=6oC) and MATs that varied between ~8.5 and 3.7oC, with a lapse rate of 0.6oC/100 m (Fig. 5-8A).
Initial model runs at an elevation of 1550±250 m (MAT=2.5±1.5oC, Ta=6oC) showed that cracking is likely to occur, but these elevations were not coincident with modeled maximum cracking intensities. Scree slopes form as rockfall deposits at the base of rock headwalls [Rapp, 1960], and so do not directly coincide with the location of maximum frost action. Scree slopes in the Southern Alps are typically 200-300 m in length [Hales and Roering, 2005], suggesting that the location of rock headwalls is a few hundred meters above the average scree slope elevation. To account for this observation, we recalculated our model with a mean elevation of 1800±250 m
(MAT=1±1.5oC, Ta=6oC) (Fig. 5-8B). In this case, the maximum predicted segregation ice depths range from 27 cm at an elevation of 1550 m to 219 cm at an elevation of 2000 m. The intensity depth profiles show that the peak cracking intensity is felt at the surface and peak temperature gradients range from 1.95 to
0.8oC/cm (Fig. 5-8B). As such, the mean rock headwall elevation (1800 m)
107
20 Ta=6oC (A)
15 762m MAT=8.46oC
C) 10
o 1554m MAT=3.71oC 5
0 Temperature (
-5 Frost Cracking Window
-10 0 60 120 180 240 300 360 Julian Day
Cracking Intensity (oC/cm) (B) 0 0.4 0.8 1.2 1.6 2 0 1550 m
1800 m 200 2000 m Depth (cm) 400
600
Figure 5-8. Yearly atmospheric temperature variation in the Southern Alps (A) and the predicted depth of segregation ice growth determined by our model (B). (A) Annual air temperatures from the Cragieburn Range (1554 m) and Mount Cook (762 m). (B) The depth of segregation ice growth at an elevation of 1800±250 m in the Southern Alps (MAT=1±1.5oC, Ta=6oC). 108
coincided with the most intense cracking predicted by our model (between 2 and
0oC). We predict elevations above 2000 m (or below the 0oC isotherm) to have considerably lower frost cracking intensities [coincident with the “high and frozen” condition ofHales and Roering, 2005], while elevations below 1600 m will have low frost cracking intensities.
We also compared the depth penetration of segregation ice growth with measures of rock strength and fracture spacing in the Southern Alps. The Torlesse
Supergroup rocks that are pervasive (sandstones and argillites) are weak and highly fractured, with a mean joint spacing <20 cm [Augustinus, 1995b; MacKinnon, 1983].
Our model results predict a maximum cracking depth of 27-219 cm and there is a high likelihood of fractures intersecting the zone of segregation ice growth.
5.7.3. Utah Rockfall Inventory
Lastly, we compared the elevation of predicted maximum segregation ice growth to a rockfall inventory dataset collected by the Utah Department of
Transportation [Pack and Boie, 2002]. This dataset was gathered to understand road hazards from rockfall. Rockfall frequency (expressed as rockfalls/year) and block size data were combined with measures of rockfall susceptibility (e.g. the presence of geotechnical structures that reduce amount of road failures) to create a rockfall hazard rating [Pack and Boie, 2002]. We focused on the relationship between rockfall frequency and elevation. Rockfall frequency was calculated by recording of the 109
3500 (A) -4 -2 3000 0 2500 2
2000 4 C) o tion (m) tion
a 6
1500 MAT (
Elev 8 1000 10
500 12 14 0
0 40 80 120 160 Rockfall Frequency (falls/year)
Callao 25 (B) 1316 m, MAT = 7.2oC 20
15 Coalville 1945 m, MAT= 5.1oC C) o 10
5 Electric Lake 2540 m, MAT=1.2oC 0 Temperature (
-5 Frost Cracking Window -10
-15 0 90 180 270 360 Julian Day
Figure 5-9. Elevation and frequency of rockfall events on roads in Utah [Pack and Boie, 2002]. (A) There is a distinctive peak in rockfall frequency above 2000 meters elevation reflecting the first influence of segregation ice growth. (B) Annual temperature curves for three temperature stations in Utah (Callao, Coalville, and Electric Lake) that were used to create predictions of frost cracking depths used for comparison with the rockfall frequency data. 110
number of times a road crew had to collect debris from a section of road. This provides a sample of rock headwalls adjacent to roads. Elevations were measured at the road with a median hillslope length of 156 m (with and interquartile range of 96-
283 m). Rockfall frequency varies between 0 and 160 falls/year, with a peak in rockfall intensity at road elevations between 2100 and 2800 m (Fig. 5-9A).
Because this dataset was collected across the state of Utah, it covers a range of rock types and climatic regimes. Nonetheless, high mountains are mostly found in the center and northwest sections of the state. We estimated the relationship between elevation and frost penetration adopting a similar technique to that used for the
Southern Alps dataset, first comparing atmospheric temperature measurements from
10 sites in the northwest and center of the state (representative temperature data are presented in Fig 5-9B) to attain a lapse rate (0.55oC/100m). The annual temperature variation at most Utah sites is ~12oC. The zone of maximum rockfall frequency corresponds to MATs between 3.5 and -0.5oC. Model results for these MATs and a
Ta of 12oC show that 400 cm is the deepest to which frost action will penetrate and exploit any fractures at shallower depths, given high cracking intensities (Fig 5-7C).
Our model predicts that the most intense cracking occurs at elevations just above the
0oC isotherm. The coincidence of the most the highest rockfall frequencies and temperatures close to the 0oC isotherm supports a segregation ice origin for the high frequencies. At higher elevations (lower MAT) frost action may penetrate deeper, but the maximum cumulative gradient is approximately an order of magnitude less. The 111
number of roads decreases at elevations above 3000 m and the small sample size may also be the cause of the drop in rockfall frequencies at these elevations. Given the heterogeneity of rock types and fracture spacing across Utah, the coincidence of the most intense predicted cracking intensity and the highest rockfall frequencies, suggests that the cumulative temperature gradient is important. Unlike the relatively simple New Zealand example, rock type, rock fracture spacing, earthquake frequency, and time since deglaciation, vary considerably across the state. Despite these complications, a first-order relationship appears to exist between temperature and rockfall frequency.
5.7.4. Implications for Mountain Range Evolution
We have presented a predictive model for the production of segregation ice in rock masses. For a given mountain range, our model shows that segregation ice weathering may be concentrated in a small elevation range. Having intense weathering concentrated within a small elevation range may affect the relief and absolute elevation for mountains with rock types that are susceptible to segregation ice induced rockfall erosion. For mountains with weak, highly fractured rock masses
(e.g. Southern Alps, New Zealand (Fig. 5-1)), peak elevations may be coincident with the frost cracking window. In contrast, areas with strong, coherent rock masses (e.g.
Sierra Nevada Mountains, California) may be less susceptible segregation ice weathering, allowing them to maintain higher peak heights. In the Southern Alps, 112
frost-driven rockfall erosion dominates within a narrow elevation band between 1550 and 2000 m [Hales and Roering, 2005] and less than 0.5% of the area of the range is greater than 2000 m. Conversely, the peak elevations of the Sierra Nevada (Mt.
Whitney, 4421 m) are approximately 1000 m higher than the 0.001oC isotherm, where frost cracking is most efficient. Mt Whitney granodiorite has a large fracture spacing [Stock et al., 2006] and is likely less affected by segregation ice weathering.
An interesting result of our modeling is the dramatic change of behavior across the 0oC isotherm, from rapid segregation ice growth close to the surface to a more diffuse ice growth at depth. In our model, this dramatic change in behavior is consistent with a change in water availability in the subsurface (the result of a 0oC condition for frost cracking). We suggest that this transition is important in mountainous landscapes as it controls the upper boundary of the frost cracking window described above. At elevations with a MAT above 0oC, only seasonal ice exists, allowing for water to be drawn readily from both the surface and groundwater.
As a result, segregation ice growth is rapid and efficient, generating high rates of frost cracking. At MATs below 0oC, segregation ice growth rates are an order of magnitude slower, resulting in a high and frozen condition.
Our prediction of an elevation-dependent zone of segregation ice weathering has implications for rockfall erosion on glacial-interglacial timescales. Our model predicts that the elevation range of effective frost processes will be affected by major changes in MAT caused by global warming and cooling events. For high mountains, 113
such as the Himalaya, a warming climate will push the zone of maximum frost cracking up in elevation, potentially affecting the amount of rockfall in high mountain passes. For low, vegetated mountains, such as the Appalachians, a cooling climate may lower the zone of maximum frost cracking erosion, potentially increasing the rate of mechanical weathering. A cooling climate is commensurate with expansion of glaciers and apparently enhanced glacial erosion [called a glacial "buzzsaw";
Brozovic et al., 1997]. Our model suggests that the elevation of the frost cracking zone is lowered in concert with the expansion of ice masses. Thus global cooling may increase the intensity of segregation ice weathering in glaciated areas and aid in the rapid erosion of mountain peaks (a rockfall “buzzsaw”?).
5.8. Conclusions
The rate of periglacial erosion in mountainous landscapes is a largely unknown quantity, with much debate existing about its role relative to fluvial, glacial and hillslope processes. We have attempted to quantify the role of ice-driven erosion through the use of a simple one dimensional heat flow model. We use this model to predict the depth and intensity of segregation ice growth in rocks using mean annual atmospheric temperature and annual temperature variations. We predict that segregation ice will grow in this model when three simple criteria are met: (1) the rock temperature is between -3 and -8oC, (2) one of the boundaries has a temperature above 0oC, meaning that water is freely available to the system, and (3) the water is 114
drawn from warmer to colder parts of the system. The growth rate of segregation ice in this system is dependent on the temperature gradient. We show that intense segregation ice growth occurs close to the surface of the rock mass for MATs above
0oC, while at negative MATs segregation ice forms at depths below 50 cm from the surface. Changes in MAT are manifest by changes in the rate and depth of segregation ice growth, while changes in Ta primarily affect the depth of ice penetration. For areas of seasonal ice growth (temperatures above 0oC), intense frost action occurs close to the surface during the autumn and winter, while weaker segregation ice growth in the presence of permafrost occurs mostly in the spring. We suggest that our predictions of the depth and intensity of segregation ice growth can be related to the fracture spacing of a rock mass to predict the susceptibility of a mountain range to frost action. Examples from the Southern Alps and Utah show that predicted maxima in the intensity of segregation ice growth coincide with maxima in rockfall erosion rates. 115
BIBLIOGRAPHY
Chapter I
Adams, J. (1980), Contemporary uplift and erosion of the Southern Alps, New Zealand, Geological Society of America Bulletin, 91, 1-114.
Batt, G.E., and J. Braun (1999), The tectonic evolution of the Southern Alps, New Zealand: insights from fully thermally coupled dynamic modelling, Geophysical Journal International, 136, 403-420.
Beaumont, C., P.J.J. Kamp, J. Hamilton, and P. Fullsack (1996), The continental collision zone, South Island, New Zealand: Comparison of geodyamical models and observations, Journal of Geophysical Research, 101 (B2), 3333- 3359.
Brandon, M.T., M.K. Roden-Tice, and J.I. Garver (1998), Late Cenozoic exhumation of the Cascadia accretionary wedge in the Olympic Mountains, northwest Washington State, Geological Society of America Bulletin, 110 (8), 985-1009.
Bull, W.B., and P.L.K. Knuepfer (1987), Adjustments by the Charwell River, New Zealand, to Uplift and Climatic Changes, Geomorphology, 1, 15-32.
Darwin, C. (1859), On the origin of species by means of natural selection, or the preservation of favoured races in the struggle for life., 672 pp., John Murray, Albermarle St., London.
Davis, W.M. (1899), The Geographical Cycle, Geographical Journal, 14, 481-504.
Ducea, M.N., and J.B. Saleeby (1996), Buoyancy sources for a large, unrooted mountain range, the Sierra Nevada, California: Evidence from xenolith thermobarometry, Journal of Geophysical Research, 101 (B4), 8229-8244.
Gilbert, G.K. (1877), Geology of the Henry Mountains (Utah), 172 pp., United States Government Printing Office, Washington D.C.
Griffiths, G.A. (1979), High sediment yields from major rivers of the Western Southern Alps, New Zealand, Nature, 282, 61-63.
116
Griffiths, G.A. (1981), Some suspended sediment yields from South Island catchments, New Zealand, Water Resources Bulletin, 17 (4), 662-671.
Hack, J.T. (1960), Interpretation of Erosional Topography in Humid Temperate Regions, American Journal of Science, 258-A, 80-97.
Hicks, D.M., M.J. McSaveney, and T.J.H. Chinn (1990), Sedimentation in Proglacial Ivory Lake, Southern Alps, New Zealand, Arctic & Alpine Research, 22 (1), 26-42.
Hovius, N., C.P. Stark, and P.A. Allen (1997), Sediment flux from a mountain belt derived by landslide mapping, Geology, 25 (3), 231-234.
Koons, P.O. (1990), Two-sided orogen: Collision and erosion from the sandbox to the Southern Alps, New Zealand, Geology, 18, 679-682.
Koons, P.O. (1994), Three-dimensional critical wedges: Tectonics and topography in oblique collisional orogens, Journal of Geophysical Research, 99 (B6), 12301-12315.
Korup, O. (2005), Large landslides and their effect on sediment flux in South Westland, New Zealand, Earth Surface Processes and Landforms, 30, 305- 307.
Molnar, P., and P. England (1990), Late Cenozoic uplift of mountain ranges and global climate change: chicken or egg?, Nature, 346, 29-34.
Norris, R.J., and A.F. Cooper (2000), Late Quaternary slip rates and slip partitioning on the Alpine Fault, New Zealand, Journal of Structural Geology, 23, 507- 520.
Raymo, M.E., and W.F. Ruddiman (1992), Tectonic forcing of late Cenozoic climate, Nature, 359, 117-122.
Reneau, S.L., and W.E. Dietrich (1991), Erosion rates in the Southern Oregon Coast Range: Evidence for an equilibrium between hillslope erosion and sediment yield, Earth Surface Processes and Landforms, 16, 307-322.
Saleeby, J.B., and Z. Foster (2004), Topographic response to mantle lithosphere removal in the southern Sierra Nevada region, California, Geology, 32 (3), 245-248. 117
Tippett, J.M., and P.J.J. Kamp (1993), Fission track analysis of the late Cenozoic vertical kinematics of Continental Pacific Crust, South Island, New Zealand, Journal of Geophysical Research, 98 (B9), 16119-16148.
Wellman, H.W. (1979), An uplift map for the South Island of New Zealand, and a model for the uplift of the Southern Alps, Bulletin of the Royal Society of New Zealand, 18, 13-20.
Whipple, K.X., and B.J. Meade (2004), Controls on the strength of coupling among climate, erosion and deformation in two-sided, frictional orogenic wedges at steady state, Journal of Geophysical Research, 190 doi:10.1029/2003JF000019.
Whitehouse, I.E. (1988), Geomorphology of the central Southern Alps, New Zealand: the interaction of plate collision and atmospheric circulation, Zeitschrift fur Geomorphologie Supplementeband, 69, 105-116.
Whitehouse, I.E., and M.J. McSaveney (1983), Diachronous talus surfaces in the Southern Alps, New Zealand, and their implications to talus accumulation, Arctic and Alpine Research, 15, 53-64.
Willett, S.D., and M.T. Brandon (2002), On steady states in mountain belts, Geology, 30 (2), 175-178.
Willett, S.D., R.L. Slingerland, and N. Hovius (2001), Uplift, shortening, and steady state topography in active mountain belts, American Journal of Science, 301, 455-485.
Wunsch, C. (2004), Quantiative estimate of the Milankovitch-forced contribution to observed Quaternary climate change, Quaternary Science Reviews, 23, 1001- 1012.
Chapter II
Anderson, D.L. (1994a), The sublithospheric mantle as the source of continental flood basalts; the case against the continental lithosphere and plume head reservoirs, Earth and Planetary Science Letters, 123, 269-280. 118
Anderson, R.S. (1994b), Evolution of the Santa Cruz Mountains, California, through tectonic growth and geomorphic decay, Journal of Geophysical Research, 99 (B10), 20161-20179.
Camp, V.E. (1981), Geologic studies of the Columbia Plateau: Part II. Upper Miocene basalt distribution, reflecting source locations, tectonism, and drainage history in the Clearwater embayment Idaho, Geological Society of America Bulletin, 92, 669-678.
Camp, V.E., and P.R. Hooper (1981), Geologic studies of the Columbia Plateau; Part 1, Late Cenozoic evolution of the southeastern part of the Columbia River basalt province, Geological Society of America Bulletin, 92, 659-668.
Camp, V.E., and M.E. Ross (2004), Mantle dynamics and genesis of mafic magmatism in the intermontane Pacific Northwest, Journal of Geophysical Research, 109 (B08204), doi:10.1029/2003JB002838.
Camp, V.E., M.E. Ross, and W.E. Hanson (2003), Genesis of flood basalts and Basin and Range volcanic rocks from Steens Mountain to the Malheur River Gorge, Oregon, Geological Society of America Bulletin, 115 (1), 105-128.
Dodson, A., M.B. Kennedy, and D.J. De Paolo (1997), Helium and neon isotopes in the Imnaha Basalt, CRBG, evidence for a Yellowstone plume source, Earth and Planetary Science Letters, 150 (3-4), 1997.
Elkins-Tanton, L.T., and B.H. Hager (2000), Melt intrusion as a trigger for lithospheric foundering and the eruption of the Siberian flood basalts, Geophysical Research Letters, 27 (23), 3937-3940.
Goles, G.G. (1986), Miocene Basalts of the Blue Mountains Province in Oregon. I: Compositional Types and their Geological Settings, Journal of Petrology, 27 (2), 495-520.
Hammond, W.C., and E.D. Humphreys (2000), Upper mantle seismic wave velocity: The effect of realistic partial melt geometries, Journal of Geophysical Research, 105, 10975-10999.
Hooper, P.R. (1997), The Columbia River Flood Basalt Province: Current Status, Geophysical Monograph, 100, 1-27.
Hooper, P.R., and V.E. Camp (1981), Deformation of the southeast part of the Columbia Plateau, Geology, 9, 323-328. 119
Hooper, P.R., and R.M. Conrey (1989), A model for the tectonic setting of the Columbia River basalt eruptions, Geological Society of America Special Paper, 239, 293-306.
Hooper, P.R., G.D. Webster, and V.E. Camp (1985), Geologic Map of the Clarkson 15 Minute Quadrangle, Idaho and Washington, State of Washington Department of Natural Resources Geologic Map Series, GM-31.
Jordan, B.T. (2005), The Oregon High Lava Plains: A province of counter-tectonic age-progressive volcanism, in Plumes, Plates, & Paradigms, edited by G.R. Foulger, J.H. Natland, D.C. Presnall, and D.L. Anderson, pp. in press.
Lallemant, H.G.A. (1994), Pre-Cretaceous Tectonic Evolution of the Blue Mountains Province, Northeastern Oregon, USGS Professional Paper, 1438, 271-304.
Lambeck, K. (1988), Geophysical Geodesy, the Slow Deformations of the Earth, 718 pp., Clarendon Press, Oxford.
McCaffrey, R., M.D. Long, C. Goldfinger, P.C. Zwick, J.L. Nablek, C.K. Johnson, and C. Smith (2000), Rotation and plate locking at the southern Cascadia subduction zone, Geophysical Research Letters, 27 (19), 3117-3120.
Reidel, S.P., P.R. Hooper, G.D. Webster, and V.E. Camp (1992), Geologic map of Southeastern Asotin County, Washington, Washington Division of Geology and Earth Resources.
Reidel, S.P., T.L. Tolan, P.R. Hooper, M.V. Beeson, K.R. Fecht, R.D. Bentley, and J.L. Anderson (1989), The Grande Ronde Basalt, Columbia River Basalt Group; Stratigraphic descriptions and correlations in Washington, Oregon, and Idaho, Geological Society of America Special Paper, 239, 21-54.
Saleeby, J.B., and Z. Foster (2004), Topographic response to mantle lithosphere removal in the southern Sierra Nevada region, California, Geology, 32 (3), 245-248.
Saltzer, R.L., and E.D. Humphreys (1997), Upper mantle P-wave structure of the eastern Snake River Plain and its relationship to geodynamic models of the region, Journal of Geophysical Research, 102, 11 829-11841.
Schutt, D.L., and C.E. Lesher (2005), The effect and seismic velocity of garnet and spinel lherzolite, Journal of Geophysical Research, in press. 120
Tolan, T.L., S.P. Reidel, M.V. Beeson, J.L. Anderson, K.R. Fecht, and F.J. Swanson (1989), Revisions to the estimates of the areal extent and volume of the Columbia River Basalt Group, Geological Society of America Special Paper, 239, 1-20.
Walker, G.W. (1979), Reconnaissance geologic map of the Oregon part of the Grangeville Quadrangle, Baker, Union, Umatilla, and Wallowa Counties, Oregon, USGS Miscellaneous Investigations Series, Map I-1116.
Wells, R.E., C.S. Weaver, and R.J. Blakely (1998), Fore-arc migration in Cascadia and its neotectonic significance, Geology, 26 (8), 759-762.
Chapter III
Adams, J. (1980), Contemporary uplift and erosion of the Southern Alps, New Zealand, Geological Society of America Bulletin, 91, 1-114.
Anderson, R.S. (1998), Near-surface thermal profiles in alpine bedrock: Implications for the frost weathering of rock, Arctic & Alpine Research, 30 (4), 362-372.
Augustinus, P. (1995), Glacial valley cross-profile development: the influence of in situ rock stress and rock mass strength, with examples from the Southern Alps, New Zealand, Geomorphology, 14, 87-97.
Ballantyne, C.K., and J.D. Eckford (1984), Characteristics and evolution of two relict talus slopes in Scotland, Scottish Geographical Magazine, 100 (1), 20-33.
Batt, G.E., and M.T. Brandon (2002), Lateral thinking: 2-D interpretation of thermochronolgy in convergent orogen settings, Tectonophysics, 349, 185- 201.
Brook, M.S., and J.M. Tippett (2002), The influence of rock mass strength on the form and evolution of deglaciated valley slopes in the English Lake District, Scottish Journal of Geology, 38, 15-20. 121
Brozovic, N., D.W. Burbank, and A. Meigs (1997), Climatic limits on landscape development in the northwestern Himalaya, Science, 276 (571-574).
Burbank, D.W., J. Leland, E. Fielding, R.S. Anderson, N. Brozovic, M.R. Reid, and C. Duncan (1996), Bedrock incision, rock uplift and threshold hillslopes in the northwestern Himalayas, Nature, 379, 505-510.
Deere, D.U., and D.W. Deere (1988), The Rock Quality Designation (RQD) Index in Practice, in Rock Classification Systems for Engineering Purposes, edited by L. Kirkaldie, pp. 91-101, American Society for Testing and Materials, Philidelphia.
Griffiths, G.A. (1981), Some suspended sediment yields from South Island catchments, New Zealand, Water Resources Bulletin, 17 (4), 662-671.
Griffiths, G.A., and M.J. McSaveney (1983), Distribution of mean annual precipitation across some steepland regions of New Zealand, New Zealand Journal of Science, 36, 197-209.
Hallet, B., J. Walder, and C.W. Stubbs (1991), Weathering by segregation ice growth in microcracks at sustained sub-zero temperatures: Verification from an experimental study using acoustic emissions, Permafrost and Periglacial Processes, 2, 283-300.
Hirano, M., and M. Aniya (1988), A rational explanation of cross-profile morphology for glacial valleys and of glacial valley development, Earth Surface Processes and Landforms, 13, 707-716.
Hovius, N., C.P. Stark, and P.A. Allen (1997), Sediment flux from a mountain belt derived by landslide mapping, Geology, 25 (3), 231-234.
Ives, J.D., and B.D. Fahey (1971), Permafrost occurrence in the Front Range, Colorado Rocky Mountains, U.S.A., Journal of Glaciology, 10 (58), 105-111.
122
Ivy-Ochs, S., C. Schuluchter, P.W. Kubik, and G.H. Denton (1999), Moraine exposure dates imply synchronous Younger Dryas glacier advances in the European Alps and in the Southern Alps of New Zealand, Geografiska Annaler, 81 A (2), 313-323.
MacKinnon, T.C. (1983), Origin of the Torlesse terrane and coeval rocks, South Island, New Zealand, Geological Society of America Bulletin, 94, 967-985.
Matsuoka, N. (2001), Direct observation of frost wedging in alpine bedrock, Earth Surface Processes and Landforms, 26, 601-614.
McSaveney, M.J. (2002), Recent rockfalls and rock avalanches in Mount Cook National Park, New Zealand, Reviews in Engineering Geology, 15, 35-70.
Miller, D.J., and T. Dunne (1996), Topographic perturbations of regional stresses and consequent bedrock fracturing, Journal of Geophysical Research, 101 (B11), 25523-25536.
Pettinga, J.R., M.D. Yetton, R.J. Van Dissen, and G. Downes (2001), Earthquake source identification and characterisation for the Canterbury Region, South Island, New Zealand, Bulletin of the New Zealand Society for Earthquake Engineering, 34, 282-317.
Reneau, S.L., and W.E. Dietrich (1991), Erosion rates in the Southern Oregon Coast Range: Evidence for an equilibrium between hillslope erosion and sediment yield, Earth Surface Processes and Landforms, 16, 307-322.
Sass, O., and K. Wollny (2001), Investigations regarding alpine talus slopes using ground-penetrating radar (GPR) in the Bavarian Alps, Germany, Earth Surface Processes and Landforms, 26, 1071-1086.
Schmidt, K., and D.R. Montgomery (1995), Limits to relief, Science, 270, 617-620.
123
Suggate, R.P. (1990), Late Pliocene and Quaternary glaciations of New Zealand, Quaternary Science Reviews, 9, 175-197.
Tippett, J.M., and P.J.J. Kamp (1993), Fission track analysis of the late Cenozoic vertical kinematics of Continental Pacific Crust, South Island, New Zealand, Journal of Geophysical Research, 98 (B9), 16119-16148.
Walcott, R.I. (1998), Modes of Oblique Compression: Late Cenozoic tectonics of the South Island of New Zealand, Reviews of Geophysics, 36 (1), 1-26.
Walder, J., and B. Hallet (1985), A theoretical model of the fracture of rock during freezing, Geological Society of America Bulletin, 96, 336-346.
Whipple, K.X., E. Kirby, and S.H. Brocklehurst (1999), Geomorphic limits to climate-induced increases in topographic relief, Nature, 401, 39-43.
Whitehouse, I.E. (1988), Geomorphology of the central Southern Alps, New Zealand: the interaction of plate collision and atmospheric circulation, Zeitschrift fur Geomorphologie Supplementeband, 69, 105-116.
Wieczorek, G.F. (2002), Catastrophic rockfalls and rockslides in the Sierra Nevada, USA, Reviews in Engineering Geology, 15, 165-190.
124
Chapter IV
Adams, C.J.D. (1981), Uplift rates and thermal structure in the Alpine fault zone and Alpine Schist, Southern Alps, New Zealand, in Thrust and Nappe Tectonics, edited by K. McClay, and N.J. Price, pp. 211-222, Geological Society [London] Special Publication.
Adams, J. (1980), Contemporary uplift and erosion of the Southern Alps, New Zealand, Geological Society of America Bulletin, 91, 1-114.
Almond, P.C., N.T. Moar, and O.B. Lian (2001), Reinterpretation of the glacial chronology of South Westland, New Zealand, New Zealand Journal of Geology and Geophysics, 44, 1-15.
Anderson, B., and A. Mackintosh (2006), Temperature change is the major driver of late-glacial and Holocene glacier fluctuations in New Zealand, Geology, 34 (2), 121-124.
Anderson, R.S. (1998), Near-surface thermal profiles in alpine bedrock: Implications for the frost weathering of rock, Arctic & Alpine Research, 30 (4), 362-372.
Arthur, G.R. (1975), Late Pleistocene-Holocene geomorphology and neotectonics of part of the Waimakariri Valley, Undergraduate thesis, University of Centerbury, Christchurch.
Augustinus, P. (1992), The influence of rock mass strength on glacial valley cross profile morphometry: A case study from the Southern Alps, New Zealand, Earth Surface Processes and Landforms, 17, 39-51.
Augustinus, P. (1995a), Glacial valley cross-profile development: the influence of in situ rock stress and rock mass strength, with examples from the Southern Alps, New Zealand, Geomorphology, 14, 87-97.
Augustinus, P. (1995b), Rock mass strength and the stability of some glacial valley slopes, Zeitschrift fur Geomorphologie, 39 (1), 55-68.
Ballantyne, C.K. (2002), A general model of paraglacial landscape response, The Holocene, 12, 371-376.
Ballantyne, C.K., and J.D. Eckford (1984), Characteristics and evolution of two relict talus slopes in Scotland, Scottish Geographical Magazine, 100 (1), 20-33. 125
Ballantyne, C.K., and M.P. Kirkbride (1987), Rockfall activity in upland Britain during the Loch Lomond stadial, The Geographical Journal, 153 (1), 86-92.
Basher, L.R., P. Tonkin, and M.J. McSaveney (1988), Geomorphic history of a rapidly uplifting area on a compressional plate boundary: Cropp River, New Zealand, Zeitschrift fur Geomorphologie Supplementeband, 69, 117-131.
Batt, G.E. (2001), The approach to steady-state theromochronological distribution following orogenic development in the Southern Alps of New Zealand, American Journal of Science, 301, 374-384.
Batt, G.E., and M.T. Brandon (2002), Lateral thinking: 2-D interpretation of thermochronolgy in convergent orogen settings, Tectonophysics, 349, 185- 201.
Batt, G.E., and J. Braun (1999), The tectonic evolution of the Southern Alps, New Zealand: insights from fully thermally coupled dynamic modelling, Geophysical Journal International, 136, 403-420.
Batt, G.E., J. Braun, B.P. Kohn, and I. McDougall (2000), Thermochronological analysis of the dynamics of the Southern Alps, New Zealand, Geological Society of America Bulletin, 112 (2), 250-266.
Beavan, J., M. Denham, P. Denys, B.H. Hager, T. Herring, C. Kurnik, D. Matheson, P. Molnar, and C. Pearson (2002), A direct measurement of the uplift rate of the Southern Alps, Eos, Transactions, American Geophysical Union, 83, SE31D-01.
Benn, J.L. (2005), Landslide events on the West Coast, South Island, 1867-2002, New Zealand Geographer, 61, 3-13.
Brocklehurst, S.H., and K.X. Whipple (2004), Hypsometry of glaciated landscapes, Earth Surface Processes and Landforms, 29, 907-926.
Brook, M.S., M.P. Kirkbride, and B.W. Brock (2006), Quantified time scale for glacial valley cross-profile evolution in mountains, Geology, 34 (8), 637-640.
Bull, W.B., and M.T. Brandon (1998), Lichen dating of earthquake generated regional rockfall events, Southern Alps, New Zealand, Geological Society of America Bulletin, 110 (1), 60-84. 126
Bull, W.B., and A.F. Cooper (1986), Uplifted marine terraces along the Alpine Fault, New Zealand, Science, 234, 1225-1228.
Burrows, C.J. (1975), A 500-year old landslide in the Acheron River Valley, Canterbury, New Zealand Journal of Geology and Geophysics, 18 (2), 357- 360.
Carter, R.M. (2005), A New Zealand climatic template back to c.3.9 Ma: ODP Site 1119, Canterbury Bight, south-west Pacific Ocean, and its relationship to onland sucessions, Journal of the Royal Society of New Zealand, 35, 9-42.
Chamberlain, C.P., P.K. Zeitler, and A.F. Cooper (1995), Geochronologic constraints of the uplift and metamorphism along the Alpine fault, South Island, New Zealand, New Zealand Journal of Geology and Geophysics, 38, 515-523.
Chinn, T.J.H. (1981), Use of rock-weathering rind thickness for Holocene absolute age-dating in New Zealand, Arctic & Alpine Research, 13, 33-45.
Cockayne, L. (1958), The Vegetation of New Zealand, 456 pp., Hafner Publishing Co., New York.
Cooper, A.F., and D.G. Bishop (1979), Uplift rates and high level marine platforms associated with the Alpine Fault at Okuru River, South Westland, Bulletin of the Royal Society of New Zealand, 18, 35-43.
Coutard, J.-P., and B. Francou (1989), Rock temperature measurements in two alpine environments: Implications for frost shattering, Arctic & Alpine Research, 21 (4), 399-416.
Deere, D.U., and D.W. Deere (1988), The rock quality designation (RQD) index in practice, in Rock Classification Systems for Engineering Purposes, edited by L. Kirkaldie, pp. 91-101, American Society for Testing and Materials, Philidelphia.
Edgar, E., and H.E. Conner (1999), Species novae graminum Novae-Zelandiae I, New Zealand Journal of Botany, 37, 63-70.
Fisher, F.J.F. (1952), Observations on the vegetation of screes in Canterbury, New Zealand, Journal of Ecology, 40, 156-167.
127
Gage, M. (1958), Late Pleistocene glaciations of the Waimakariri Valley, Canterbury, New Zealand, New Zealand Journal of Geology and Geophysics, 1, 123-155.
Gage, M., and R.P. Suggate (1958), Glacial chronology of the New Zealand Pleistocene, Geological Society of America Bulletin, 69, 589-598.
Grab, S.W., C.K. Gatebe, and A.M. Kinyua (2004), Ground thermal profiles from Mount Kenya, East Africa, Geografiska Annaler, 86A, 131-141.
Grapes, R. (1995), Uplift and exhumation of the Alpine schist, Southern Alps, New Zealand: themobarometric constraints, New Zealand Journal of Geology and Geophysics, 38, 525-533.
Griffiths, G.A. (1979), High sediment yields from major rivers of the Western Southern Alps, New Zealand, Nature, 282, 61-63.
Griffiths, G.A. (1981), Some suspended sediment yields from South Island catchments, New Zealand, Water Resources Bulletin, 17 (4), 662-671.
Griffiths, G.A., and M.J. McSaveney (1983a), Distribution of mean annual precipitation across some steepland regions of New Zealand, New Zealand Journal of Science, 36, 197-209.
Griffiths, G.A., and M.J. McSaveney (1983b), Hydrology of a basin with extreme rainfalls-Cropp River, New Zealand, New Zealand Journal of Science, 26, 293-306.
Griffiths, G.A., and M.J. McSaveney (1986), Sedimentation and river containment on Waitangitaona alluvial fan- South Westland, New Zealand, Zeitschrift fur Geomorphologie, 30 (2), 215-230.
Hales, T.C., and J.J. Roering (2005), Climate-controlled variations in scree production, Southern Alps, New Zealand, Geology, 33 (9), 701-704.
Hales, T.C., and J.J. Roering (2006), Climatic controls on frost cracking and implications for the evolution of bedrock landscapes, Journal of Geophysical Research, in press.
Hall, K., and M.-F. Andre (2001), New insights into rock weathering from high- frequency rock temperature data: an Antarctic study of weathering by thermal stress, Geomorphology, 41, 23-35. 128
Hall, K., B.S. Lindgren, and P. Jackson (2005), Rock albedo and monitoring of thermal conditions in respect of weathering: some expected and some unexpected results, Earth Surface Processes and Landforms, 30, 801-811.
Hallet, B., J. Walder, and C.W. Stubbs (1991), Weathering by segregation ice growth in microcracks at sustained sub-zero temperatures: Verification from an experimental study using acoustic emissions, Permafrost and Periglacial Processes, 2, 283-300.
Halsey, D.P., D.J. Mitchell, and S.J. Dews (1998), Influence of climatically induced cycles in physical weathering, Quarterly Journal of Engineering Geology, 31, 359-367.
Harris, S.A. (1975), Petrology and origin of stratified scree in New Zealand, Quaternary Research, 5, 199-214.
Henderson, R.D., and S.M. Thompson (1999), Extreme rainfalls in the Southern Alps of New Zealand, Journal of Hydrology (New Zealand), 38 (2), 309-330.
Hicks, D.M., M.J. McSaveney, and T.J.H. Chinn (1990), Sedimentation in proglacial Ivory Lake, Southern Alps, New Zealand, Arctic & Alpine Research, 22 (1), 26-42.
Hinchliffe, S., and C.K. Ballantyne (1999), Talus accumulation and rockwall retreat, Trotternish, ISle of Skye, Scotland, Scottish Geographical Journal, 115 (1), 53-70.
Hovius, N., C.P. Stark, and P.A. Allen (1997), Sediment flux from a mountain belt derived by landslide mapping, Geology, 25 (3), 231-234.
Iqbal, M. (1983), An Introduction to Solar Radiation, 390 pp., Academic Press, Toronto.
Ishikawa, M., Y. Kurashige, and K. Hirakawa (2004), Analysis of crack movements observed in an alpine bedrock cliff, Earth Surface Processes and Landforms, 29, 883-891.
Ivy-Ochs, S., C. Schuluchter, P.W. Kubik, and G.H. Denton (1999), Moraine exposure dates imply synchronous Younger Dryas glacier advances in the 129
European Alps and in the Southern Alps of New Zealand, Geografiska Annaler, 81 A (2), 313-323.
Jacobson, A.D., and J.D. Blum (2003), Relationship between mechanical erosion and atmospheric COsub2 consumption in the New Zealand Southern Alps, Geology, 31 (10), 865-868.
Kamp, P.J.J., P.F. Green, and S. White (1989), Fission track analysis reveals character of collisional tectonics in New Zealand, Tectonics, 8 (2), 169-195.
Kirkbride, M.P., and D. Matthews (1997), The role of fluvial and glacial erosion in landscape evolution: The Ben Ohau Range, New Zealand, Earth Surface Processes and Landforms, 22, 317-327.
Koons, P.O. (1994), Three-dimensional critical wedges: Tectonics and topography in oblique collisional orogens, Journal of Geophysical Research, 99 (B6), 12301-12315.
Korup, O. (2004a), Geomorphic implications of fault zone weakening: slope instability along the Alpine Fault, South Westland to Fiordland, New Zealand Journal of Geology and Geophysics, 47, 257-267.
Korup, O. (2004b), Landslide-induced river channel avulsions in mountain catchments of southwest New Zealand, Geomorphology, 63, 57-80.
Korup, O. (2005a), Distribution of landslides in southwest New Zealand, Landslides, 2, 43-51.
Korup, O. (2005b), Geomorphic imprint of landslides on alpine river systems, southwest New Zealand, Earth Surface Processes and Landforms, 30, 783- 800.
Korup, O. (2005c), Large landslides and their effect on sediment flux in South Westland, New Zealand, Earth Surface Processes and Landforms, 30, 305- 307.
Korup, O. (2006a), Effects of deep-seated landslides on hillslope morphology, western Southern Alps, New Zealand, Journal of Geophysical Research, 111 (F01018), doi: 10.1029/2004JF000242.
130
Korup, O. (2006b), Rock-slope failure and the river long profile, Geology, 34 (1), 45- 48.
Korup, O., and M. Crozier (2002), Landslide types and geomorphic impact on river channels, Southern Alps, New Zealand, in Proceedings of the First European Conference on Landslides, edited by J. Rybar, J. Stemberk, and P. Wagner, pp. 233-238, Prague.
Korup, O., M.J. McSaveney, and T.R.H. Davies (2004), Sediment generation and delivery from large historic landslides in the Southern Alps, New Zealand, Geomorphology, 61, 189-207.
Korup, O., J. Schmidt, and M.J. McSaveney (2005), Regional relief characteristics and denudation pattern of the western Southern Alps, New Zealand, Geomorphology, 71, 402-423.
Leckie, D.A. (2001), Climatic controls on nonmarine sedimentation-maximum aggradation during lowstand, incision during transgession and highstand, in American Association of Petroleum Geology Annual Meeting, pp. 114, Denver.
Little, T.A., S. Cox, J.K. Vry, and G.E. Batt (2005), Variations in exhumation level and uplift rate along the oblique-slip Alpine fault, central Southern Alps, New Zealand, Geological Society of America Bulletin, 117 (5/6), 707-723.
MacKinnon, T.C. (1983), Origin of the Torlesse terrane and coeval rocks, South Island, New Zealand, Geological Society of America Bulletin, 94, 967-985.
Matsuoka, N. (1994), Diurnal freeze-thaw depth in rockwalls:field measurements and theoretical considerations, Earth Surface Processes and Landforms, 19, 423- 435.
Matsuoka, N. (2001), Direct observation of frost wedging in alpine bedrock, Earth Surface Processes and Landforms, 26, 601-614.
Matsuoka, N., and H. Sakai (1999), Rockfall activity from an alpine cliff during thawing periods, Geomorphology, 28, 309-328.
McArthur, J.L. (1987), The characteristics, classification and origin of late Pleistocene fan deposits in the Cass Basin, Canterbury, New Zealand, Sedimentology, 34, 459-471. 131
McSaveney, M.J. (1992), A manual for weathering-rind dating of grey sandstones of the Torlesse Supergroup, New Zealand, 52 pp., Institute of Geological and Nuclear Sciences, Lower Hutt, New Zealand.
Miller, D.J., and T. Dunne (1996), Topographic perturbations of regional stresses and consequent bedrock fracturing, Journal of Geophysical Research, 101 (B11), 25523-25536.
Moon, B.P. (1984), Refinement of a technique for determining rock mass strength for geomorphological purposes, Earth Surface Processes and Landforms, 9, 189- 193.
Norris, R.J., and A.F. Cooper (2000), Late Quaternary slip rates and slip partitioning on the Alpine Fault, New Zealand, Journal of Structural Geology, 23, 507- 520.
Orwin, J.F. (1998), The application and implications of rock weathering rind dating to a large rock avalanche, Cragieburen Range, New Zealand, New Zealand Journal of Geology and Geophysics, 41, 219-223.
Porter, S.C. (1975), Equilibrium-line altitudes of late Quaternary glaciers in the Southern Alps, New Zealand, Quaternary Research, 5, 27-47.
Rapp, A. (1960), Talus slopes and mountain walls at Tempelfjorden, Spitsbergen A geomorphological study of the denudation of slopes in an arctic locality, Norsk Polarinstitutt Skr., 119 (4-7), 90-96.
Rother, H., and J. Shulmeister (2005), Synoptic climate change as a driver of late Quaternary glaciations in the mid-latitudes of the Southern Hemisphere, Climate of the Past Discussions, 1, 231-253.
Selby, M.J. (1980), A rock mass strength classification for geomorphic purposes: with tests from Antarctica and New Zealand, Zeitschrift fur Geomorphologie, 24 (1), 31-51.
Selby, M.J. (1982), Controls on the stability and inclinations of hillslopes formed on hard rock, Earth Surface Processes and Landforms, 7, 449-467.
Shulmeister, J., I. Goodwin, J. Renwick, K. Harle, L. Armand, M.S. McGlone, E. Cook, J. Dodson, P.P. Hesse, P. Mayewski, and M. Curran (2004), The 132
Southern Hemisphere westerlies in the Australasian sector over the last glacial cycle: a synthesis, Quaternary International, 118-119, 23-53.
Smith, G.M., T.R.H. Davies, M.J. McSaveney, and D.H. Bell (2006), The Acheron rock avalanche, Canterbury, New Zealand-morphology and dynamics, Landslides, 3 (1), 62-72.
Speight, R. (1916), The physiography of the Cass District, Transactions of the New Zealand Institute, 48, 145-153.
Statham, I. (1976), A scree slope rockfall model, Earth Surface Processes, 1, 43-62.
Statham, I., and S.C. Francis (1986), Influence of scree accumulation and weathering on the development of steep mountain slopes, in Hillslope Processes, pp. 245- 267, Allen & Unwin, Boston.
Suggate, R.P. (1990), Late Pliocene and Quaternary Glaciations of New Zealand, Quaternary Science Reviews, 9, 175-197.
Thompson, S.M., and J. Adams (1979), Suspended load in some major rivers of New Zealand, in Physical Hydrology-New Zealand experience, edited by D.L. Murray, and P. Ackroyd, pp. 213-228, New Zealand Hydrological Society, Wellington.
Tippett, J.M., and P.J.J. Kamp (1993), Fission track analysis of the late Cenozoic vertical kinematics of continental Pacific crust, South Island, New Zealand, Journal of Geophysical Research, 98 (B9), 16119-16148.
Tippett, J.M., and P.J.J. Kamp (1995), Geomorphic evolution of the Southern Alps, New Zealand, Earth Surface Processes and Landforms, 20, 177-192.
Tucker, G.E., and R.L. Bras (1998), Hillslope processes, drainage density, and landscape morphology, Water Resources Research, 34 (10), 2751-2764.
Walcott, R.I. (1998), Modes of oblique compression: Late Cenozoic tectonics of the South Island of New Zealand, Reviews of Geophysics, 36 (1), 1-26.
Walder, J., and B. Hallet (1985), A theoretical model of the fracture of rock during freezing, Geological Society of America Bulletin, 96, 336-346.
133
Wandres, A.M., J.D. Bradshaw, S.D. Weaver, R. Maas, T.R. Ireland, and N. Eby (2004), Provenance analysis using conglomerate clast lithologies: A case study from the Pahau Terrane of New Zealand, Sedimentary Geology.
Wardle, P. (1991), Vegetation of New Zealand, 672 pp., Cambridge University Press, Cambridge.
Wellman, H.W. (1979), An uplift map for the South Island of New Zealand, and a model for the uplift of the Southern Alps, Bulletin of the Royal Society of New Zealand, 18, 13-20.
Whitehouse, I.E. (1981), A large rock avalanche in the Cragieburn Range, Canterbury, New Zealand Journal of Geology and Geophysics, 24, 415-421.
Whitehouse, I.E. (1983), Distribution of large rock avalance deposits in the central Southern Alps, New Zealand, New Zealand Journal of Geology and Geophysics, 26, 271-279.
Whitehouse, I.E. (1988), Geomorphology of the central Southern Alps, New Zealand: the interaction of plate collision and atmospheric circulation, Zeitschrift fur Geomorphologie Supplementeband, 69, 105-116.
Whitehouse, I.E., and G.A. Griffiths (1983), Frequency and hazard of large rock avalanches in the central Southern Alps, New Zealand, Geology, 11, 331-334.
Whitehouse, I.E., and M.J. McSaveney (1983), Diachronous talus surfaces in the Southern Alps, New Zealand, and their implications to talus accumulation, Arctic and Alpine Research, 15, 53-64.
Whitehouse, I.E., M.J. McSaveney, P.L.K. Knuepfer, and T.J.H. Chinn (1986), Growth of weathering rinds on Torlesse sandstone, Southern Alps, New Zealand, in Rates of Chemical Weathering of Rocks and Minerals, edited by S.M. Colman, and D.P. Dethier, pp. 419-435, Academic Press, Orlando.
Willett, S.D. (1999), Orogeny and orography: The effects of erosion on the structure of mountain belts, Journal of Geophysical Research, 104 (B12), 28957- 28981.
Willett, S.D., and M.T. Brandon (2002), On steady states in mountain belts, Geology, 30 (2), 175-178.
134
Willett, S.D., R.L. Slingerland, and N. Hovius (2001), Uplift, shortening, and steady state topography in active mountain belts, American Journal of Science, 301, 455-485.
Chapter V
Adams, C.J.D. (1981), Uplift rates and thermal structure in the Alpine fault zone and Alpine Schist, Southern Alps, New Zealand, in Thrust and Nappe Tectonics, edited by K. McClay, and N.J. Price, pp. 211-222, Geological Society [London] Special Publication.
Adams, J. (1980), Contemporary uplift and erosion of the Southern Alps, New Zealand, Geological Society of America Bulletin, 91, 1-114.
Almond, P.C., N.T. Moar, and O.B. Lian (2001), Reinterpretation of the glacial chronology of South Westland, New Zealand, New Zealand Journal of Geology and Geophysics, 44, 1-15.
Anderson, B., and A. Mackintosh (2006), Temperature change is the major driver of late-glacial and Holocene glacier fluctuations in New Zealand, Geology, 34 (2), 121-124.
Anderson, D.L. (1994a), The sublithospheric mantle as the source of continental flood basalts; the case against the continental lithosphere and plume head reservoirs, Earth and Planetary Science Letters, 123, 269-280.
Anderson, R.S. (1994b), Evolution of the Santa Cruz Mountains, California, through tectonic growth and geomorphic decay, Journal of Geophysical Research, 99 (B10), 20161-20179.
Anderson, R.S. (1998), Near-surface thermal profiles in alpine bedrock: Implications for the frost weathering of rock, Arctic & Alpine Research, 30 (4), 362-372.
Andre, M.-F. (1997), Holocene rockwall retreat in Svalbard: A triple-rate evolution, Earth Surface Processes and Landforms, 22, 423-440.
Arthur, G.R. (1975), Late Pleistocene-Holocene geomorphology and neotectonics of part of the Waimakariri Valley, Undergraduate thesis, University of Centerbury, Christchurch. 135
Augustinus, P. (1992), The influence of rock mass strength on glacial valley cross profile morphometry: A Case Study from the Southern Alps, New Zealand, Earth Surface Processes and Landforms, 17, 39-51.
Augustinus, P. (1995a), Glacial valley cross-profile development: the influence of in situ rock stress and rock mass strength, with examples from the Southern Alps, New Zealand, Geomorphology, 14, 87-97.
Augustinus, P. (1995b), Rock mass strength and the stability of some glacial valley slopes, Zeitschrift fur Geomorphologie, 39 (1), 55-68.
Ballantyne, C.K. (2002), A general model of paraglacial landscape response, The Holocene, 12, 371-376.
Ballantyne, C.K., and J.D. Eckford (1984), Characteristics and evolution of two relict talus slopes in Scotland, Scottish Geographical Magazine, 100 (1), 20-33.
Ballantyne, C.K., and M.P. Kirkbride (1987), Rockfall activity in upland Britain during the Loch Lomond stadial, The Geographical Journal, 153 (1), 86-92.
Basher, L.R., P. Tonkin, and M.J. McSaveney (1988), Geomorphic history of a rapidly uplifting area on a compressional plate boundary: Cropp River, New Zealand, Zeitschrift fur Geomorphologie Supplementeband, 69, 117-131.
Batt, G.E. (2001), The approach to steady-state theromochronological distribution following orogenic development in the Southern Alps of New Zealand, American Journal of Science, 301, 374-384.
Batt, G.E., and M.T. Brandon (2002), Lateral thinking: 2-D interpretation of thermochronolgy in convergent orogen settings, Tectonophysics, 349, 185- 201.
Batt, G.E., and J. Braun (1999), The tectonic evolution of the Southern Alps, New Zealand: insights from fully thermally coupled dynamic modelling, Geophysical Journal International, 136, 403-420.
Batt, G.E., J. Braun, B.P. Kohn, and I. McDougall (2000), Thermochronological analysis of the dynamics of the Southern Alps, New Zealand, Geological Society of America Bulletin, 112 (2), 250-266.
Beaumont, C., P.J.J. Kamp, J. Hamilton, and P. Fullsack (1996), The continental collision zone, South Island, New Zealand: Comparison of geodyamical 136
models and observations, Journal of Geophysical Research, 101 (B2), 3333- 3359.
Beavan, J., M. Denham, P. Denys, B.H. Hager, T. Herring, C. Kurnik, D. Matheson, P. Molnar, and C. Pearson (2002), A direct measurement of the uplift rate of the Southern Alps, Eos, Transactions, American Geophysical Union, 83, SE31D-01.
Benn, J.L. (2005), Landslide events on the West Coast, South Island, 1867-2002, New Zealand Geographer, 61, 3-13.
Brandon, M.T., M.K. Roden-Tice, and J.I. Garver (1998), Late Cenozoic exhumation of the Cascadia accretionary wedge in the Olympic Mountains, northwest Washington State, Geological Society of America Bulletin, 110 (8), 985-1009.
Brocklehurst, S.H., and K.X. Whipple (2004), Hypsometry of glaciated landscapes, Earth Surface Processes and Landforms, 29, 907-926.
Brook, M.S., M.P. Kirkbride, and B.W. Brock (2006), Quantified time scale for glacial valley cross-profile evolution in mountains, Geology, 34 (8), 637-640.
Brook, M.S., and J.M. Tippett (2002), The influence of rock mass strength on the form and evolution of deglaciated valley slopes in the English Lake District, Scottish Journal of Geology, 38, 15-20.
Brozovic, N., D.W. Burbank, and A. Meigs (1997), Climatic limits on landscape development in the northwestern Himalaya, Science, 276 (571-574).
Bull, W.B., and M.T. Brandon (1998), Lichen dating of earthquake generated regional rockfall events, Southern Alps, New Zealand, Geological Society of America Bulletin, 110 (1), 60-84.
Bull, W.B., and A.F. Cooper (1986), Uplifted marine terraces along the Alpine Fault, New Zealand, Science, 234, 1225-1228.
Bull, W.B., and P.L.K. Knuepfer (1987), Adjustments by the Charwell River, New Zealand, to uplift and alimatic ahanges, Geomorphology, 1, 15-32.
Burbank, D.W., J. Leland, E. Fielding, R.S. Anderson, N. Brozovic, M.R. Reid, and C. Duncan (1996), Bedrock incision, rock uplift and threshold hillslopes in the northwestern Himalayas, Nature, 379, 505-510. 137
Burrows, C.J. (1975), A 500-year old landslide in the Acheron River Valley, Canterbury, New Zealand Journal of Geology and Geophysics, 18 (2), 357- 360.
Camp, V.E. (1981), Geologic studies of the Columbia Plateau: Part II. Upper Miocene basalt distribution, reflecting source locations, tectonism, and drainage history in the Clearwater embayment Idaho, Geological Society of America Bulletin, 92, 669-678.
Camp, V.E., and P.R. Hooper (1981), Geologic studies of the Columbia Plateau; Part 1, Late Cenozoic evolution of the southeastern part of the Columbia River basalt province, Geological Society of America Bulletin, 92, 659-668.
Camp, V.E., and M.E. Ross (2004), Mantle dynamics and genesis of mafic magmatism in the intermontane Pacific Northwest, Journal of Geophysical Research, 109 (B08204), doi:10.1029/2003JB002838.
Camp, V.E., M.E. Ross, and W.E. Hanson (2003), Genesis of flood basalts and Basin and Range volcanic rocks from Steens Mountain to the Malheur River Gorge, Oregon, Geological Society of America Bulletin, 115 (1), 105-128.
Carslaw, H.S., and J.C. Jaeger (1959), Conduction of Heat in Solids, 510 pp., Oxford University Press, London.
Carter, R.M. (2005), A New Zealand climatic template back to c.3.9 Ma: ODP Site 1119, Canterbury Bight, south-west Pacific Ocean, and its relationship to onland sucessions, Journal of the Royal Society of New Zealand, 35, 9-42.
Chamberlain, C.P., P.K. Zeitler, and A.F. Cooper (1995), Geochronologic constraints of the uplift and metamorphism along the Alpine fault, South Island, New Zealand, New Zealand Journal of Geology and Geophysics, 38, 515-523.
Chinn, T.J.H. (1981), Use of rock-weathering rind thickness for Holocene absolute age-dating in New Zealand, Arctic & Alpine Research, 13, 33-45.
Church, M., R.F. Stock, and J.M. Ryder (1979), Contemporary sedimentary environments on Baffin Island, N.W.T., Canada: Debris slope accumulations, Arctic & Alpine Research, 11 (4), 371-402.
Cockayne, L. (1958), The Vegetation of New Zealand, 456 pp., Hafner Publishing Co., New York. 138
Cooper, A.F., and D.G. Bishop (1979), Uplift Rates and High Level Marine Platforms Associated with the Alpine Fault at Okuru River, South Westland, Bulletin of the Royal Society of New Zealand, 18, 35-43.
Coutard, J.-P., and B. Francou (1989), Rock Temperature Measurements in Two Alpine Environments: Implications for Frost Shattering, Arctic & Alpine Research, 21 (4), 399-416.
Dash, J.G., A.W. Rempel, and J.S. Wettlaufer (2006), The physics of premelted ice and its geophysical consequences, Reviews of Modern Physics, 78 (2), in press.
DeCelles, P.G., and P.C. DeCelles (2001), Rates of shortening, propagation, underthrusting, and flexural wave migration in continental orogenic systems, Geology, 29 (2), 135-138.
Deere, D.U., and D.W. Deere (1988), The Rock Quality Designation (RQD) Index in Practice, in Rock Classification Systems for Engineering Purposes, edited by L. Kirkaldie, pp. 91-101, American Society for Testing and Materials, Philidelphia.
Dodson, A., M.B. Kennedy, and D.J. De Paolo (1997), Helium and neon isotopes in the Imnaha Basalt, CRBG, evidence for a Yellowstone plume source, Earth and Planetary Science Letters, 150 (3-4), 1997.
Ducea, M.N., and J.B. Saleeby (1996), Buoyancy sources for a large, unrooted mountain range, the Sierra Nevada, California: Evidence from xenolith thermobarometry, Journal of Geophysical Research, 101 (B4), 8229-8244.
Edgar, E., and H.E. Conner (1999), Species novae graminum Novae-Zelandiae I, New Zealand Journal of Botany, 37, 63-70.
Elkins-Tanton, L.T., and B.H. Hager (2000), Melt intrusion as a trigger for lithospheric foundering and the eruption of the Siberian flood basalts, Geophysical Research Letters, 27 (23), 3937-3940.
Fahey, B.D., and T.H. Lefebure (1988), The freeze-thaw weathering regime at a section of the Niagara Escarpment on the Bruce Peninsula, SOuthern Ontario, Canada, Earth Surface Processes and Landforms, 13, 293-304.
Fisher, F.J.F. (1952), Observations on the Vegetation of Screes in Canterbury, New Zealand, Journal of Ecology, 40, 156-167. 139
Frich, P., and E. Brandt (1985), Holocene talus accumulation rates,- and their influence on rock glacier growth. A case study from Igpik, Disko - West Greenland, Geografisk Tidsskrift, 85, 32-45.
Gage, M. (1958), Late Pleistocene Glaciations of the Waimakariri Valley, Canterbury, New Zealand, New Zealand Journal of Geology and Geophysics, 1, 123-155.
Gage, M., and R.P. Suggate (1958), Glacial Chronology of the New Zealand Pleistocene, Geological Society of America Bulletin, 69, 589-598.
Gardner, J.S. (1983), Rockfall frequency and distribution in the Highwood Pass area, Canadian Rocky Mountains, Zeitschrift fur Geomorphologie, 27 (3), 311-324.
Gilbert, G.K. (1877), Geology of the Henry Mountains (Utah), 172 pp., United States Government Printing Office, Washington D.C.
Gilpin, R.R. (1980), A Model for the prediction of ice lensing and frost heave in soils, Water Resources Research, 16, 918-930.
Gold, L.W., and A.H. Lachenbruch (1983), Thermal Conditions in Permafrost-A Review of North American Literature, in Permafrost: Fourth International Conference, pp. 3-25, National Academy Press, Fairbanks, AK.
Goles, G.G. (1986), Miocene Basalts of the Blue Mountains Province in Oregon. I: Compositional Types and their Geological Settings, Journal of Petrology, 27 (2), 495-520.
Grab, S.W., C.K. Gatebe, and A.M. Kinyua (2004), Ground thermal profiles from Mount Kenya, East Africa, Geografiska Annaler, 86A, 131-141.
Grapes, R. (1995), Uplift and exhumation of the Alpine schist, Southern Alps, New Zealand: themobarometric constraints, New Zealand Journal of Geology and Geophysics, 38, 525-533.
Gray, J.T. (1973), Geomorphic effects of avalanches and rock-falls on steep mountain slopes in the central Yukon territory, in Research in Polar and Alpine Geomorphology, edited by B.D. Fahey, and R.D. Thompson, pp. 107-117, Geo Abstracts Ltd, Norwich.
Griffiths, G.A. (1979), High sediment yields from major rivers of the Western Southern Alps, New Zealand, Nature, 282, 61-63. 140
Griffiths, G.A. (1981), Some suspended sediment yields from South Island catchments, New Zealand, Water Resources Bulletin, 17 (4), 662-671.
Griffiths, G.A., and M.J. McSaveney (1983a), Distribution of mean annual precipitation across some steepland regions of New Zealand, New Zealand Journal of Science, 36, 197-209.
Griffiths, G.A., and M.J. McSaveney (1983b), Hydrology of a basin with extreme rainfalls-Cropp River, New Zealand, New Zealand Journal of Science, 26, 293-306.
Griffiths, G.A., and M.J. McSaveney (1986), Sedimentation and river containment on Waitangitaona alluvial fan- South Westland, New Zealand, Zeitschrift fur Geomorphologie, 30 (2), 215-230.
Gruber, S., M. Hoelzle, and W. Haeberli (2004), Permafrost thaw and destailization of Alpine rock walls in the hot summer of 2003, Geophysical Research Letters, 31, L13504.
Hack, J.T. (1960), Interpretation of Erosional Topography in Humid Temperate Regions, American Journal of Science, 258-A, 80-97.
Hales, T.C., D.L. Abt, E.D. Humphreys, and J.J. Roering (2005), Lithospheric instability as an Origin for Columbia River Flood Basalts and Wallowa Mountains Uplift in NE Oregon, USA, Nature, 432, 842-845.
Hales, T.C., and J.J. Roering (2005), Climate-controlled variations in scree production, Southern Alps, New Zealand, Geology, 33 (9), 701-704.
Hales, T.C., and J.J. Roering (2006), Climatic Controls on Frost Cracking and Implications for the Evolution of Bedrock Landscapes, Journal of Geophysical Research, in press.
Hall, K., and M.-F. Andre (2001), New insights into rock weathering from high- frequency rock temperature data: an Antarctic study of weathering by thermal stress, Geomorphology, 41, 23-35.
Hall, K., B.S. Lindgren, and P. Jackson (2005), Rock albedo and monitoring of thermal conditions in respect of weathering: some expected and some unexpected results, Earth Surface Processes and Landforms, 30, 801-811. 141
Hallet, B., L. Hunter, and J. Bogen (1996), Rates of erosion and sediment evacuation by glaciers: A review of field data and their implications, Global and Planetary Change, 12, 213-255.
Hallet, B., J. Walder, and C.W. Stubbs (1991), Weathering by Segregation Ice Growth in Microcracks at Sustained Sub-zero Temperatures: Verification from an Experimental Study Using Acoustic Emissions, Permafrost and Periglacial Processes, 2, 283-300.
Halsey, D.P., D.J. Mitchell, and S.J. Dews (1998), Influence of climatically induced cycles in physical weathering, Quarterly Journal of Engineering Geology, 31, 359-367.
Hammond, W.C., and E.D. Humphreys (2000), Upper mantle seismic wave velocity: The effect of realistic partial melt geometries, Journal of Geophysical Research, 105, 10975-10999.
Harris, S.A. (1975), Petrology and Origin of Stratified Scree in New Zealand, Quaternary Research, 5, 199-214.
Harris, S.A., and D.E. Pederson (1998), Thermal Regimes Beneath Coarse Blocky Materials, Permafrost and Periglacial Processes, 9, 107-120.
Henderson, R.D., and S.M. Thompson (1999), Extreme Rainfalls in the Southern Alps of New Zealand, Journal of Hydrology (New Zealand), 38 (2), 309-330.
Henry, K.S. (2000), A Review of the Thermodynamics of Frost Heave, pp. 19, US Army Corps of Engineers.
Hicks, D.M., M.J. McSaveney, and T.J.H. Chinn (1990), Sedimentation in Proglacial Ivory Lake, Southern Alps, New Zealand, Arctic & Alpine Research, 22 (1), 26-42.
Hinchliffe, S., and C.K. Ballantyne (1999), Talus Accumulation and Rockwall Retreat, Trotternish, ISle of Skye, Scotland, Scottish Geographical Journal, 115 (1), 53-70.
Hirano, M., and M. Aniya (1988), A rational explanation of cross-profile morphology for glacial valleys and of glacial valley development, Earth Surface Processes and Landforms, 13, 707-716.
Hiroi, M., T. Mizusaki, T. Tsuneto, A. Hirai, and K. Eguchi (1989), Frost-heave phenomena of 4He on porous glasses, Physical Review B, 40 (10), 6581-6590. 142
Hoek, E. (1983), Strength of Jointed Rock Masses, Geotechnique, 33 (3), 187-223.
Hooper, P.R. (1997), The Columbia River Flood Basalt Province: Current Status, Geophysical Monograph, 100, 1-27.
Hooper, P.R., and V.E. Camp (1981), Deformation of the southeast part of the Columbia Plateau, Geology, 9, 323-328.
Hooper, P.R., and R.M. Conrey (1989), A model for the tectonic setting of the Columbia River basalt eruptions, Geological Society of America Special Paper, 239, 293-306.
Hooper, P.R., G.D. Webster, and V.E. Camp (1985), Geologic Map of the Clarkson 15 Minute Quadrangle, Idaho and Washington, State of Washington Department of Natural Resources Geologic Map Series, GM-31.
Hovius, N., C.P. Stark, and P.A. Allen (1997), Sediment flux from a mountain belt derived by landslide mapping, Geology, 25 (3), 231-234.
Iqbal, M. (1983), An Introduction to Solar Radiation, 390 pp., Academic Press, Toronto.
Ishikawa, M., Y. Kurashige, and K. Hirakawa (2004), Analysis of crack movements observed in an alpine bedrock cliff, Earth Surface Processes and Landforms, 29, 883-891.
Ives, J.D., and B.D. Fahey (1971), Permafrost occurrence in the Front Range, Colorado Rocky Mountains, U.S.A., Journal of Glaciology, 10 (58), 105-111.
Ivy-Ochs, S., C. Schuluchter, P.W. Kubik, and G.H. Denton (1999), Moraine exposure dates imply synchronous Younger Dryas glacier advances in the European Alps and in the Southern Alps of New Zealand, Geografiska Annaler, 81 A (2), 313-323.
Jordan, B.T. (2005), The Oregon High Lava Plains: A province of counter-tectonic age-progressive volcanism, in Plumes, Plates, & Paradigms, edited by G.R. Foulger, J.H. Natland, D.C. Presnall, and D.L. Anderson, pp. in press.
Kamp, P.J.J., P.F. Green, and S. White (1989), Fission Track Analysis Reveals Character of Collisional Tectonics in New Zealand, Tectonics, 8 (2), 169-195. 143
Kirkbride, M.P., and D. Matthews (1997), The role of fluvial and glacial erosion in landscape evolution: The Ben Ohau Range, New Zealand, Earth Surface Processes and Landforms, 22, 317-327.
Koons, P.O. (1990), Two-sided orogen: Collision and erosion from the sandbox to the Southern Alps, New Zealand, Geology, 18, 679-682.
Koons, P.O. (1994), Three-dimensional critical wedges: Tectonics and topography in oblique collisional orogens, Journal of Geophysical Research, 99 (B6), 12301-12315.
Korup, O. (2004a), Geomorphic implications of fault zone weakening: slope instaility along the Alpine Fault, South Westland to Fiordland, New Zealand Journal of Geology and Geophysics, 47, 257-267.
Korup, O. (2004b), Landslide-induced river channel avulsions in mountain catchments of southwest New Zealand, Geomorphology, 63, 57-80.
Korup, O. (2005a), Distribution of landslides in southwest New Zealand, Landslides, 2, 43-51.
Korup, O. (2005b), Geomorphic imprint of landslides on alpine river systems, southwest New Zealand, Earth Surface Processes and Landforms, 30, 783- 800.
Korup, O. (2005c), Large landslides and their effect on sediment flux in South Westland, New Zealand, Earth Surface Processes and Landforms, 30, 305- 307.
Korup, O. (2006a), Effects of deep-seated landslides on hillslope morphology, western Southern Alps, New Zealand, Journal of Geophysical Research, 111 (F01018), doi: 10.1029/2004JF000242.
Korup, O. (2006b), Rock-slope failure and the river long profile, Geology, 34 (1), 45- 48.
Korup, O., and M. Crozier (2002), Landslide types and geomorphic impact on river channels, Southern Alps, New Zealand, in Proceedings of the First European Conference on Landslides, edited by J. Rybar, J. Stemberk, and P. Wagner, pp. 233-238, Prague. 144
Korup, O., M.J. McSaveney, and T.R.H. Davies (2004), Sediment generation and delivery from large historic landslides in the Southern Alps, New Zealand, Geomorphology, 61, 189-207.
Korup, O., J. Schmidt, and M.J. McSaveney (2005), Regional relief characteristics and denudation pattern of the western Southern Alps, New Zealand, Geomorphology, 71, 402-423.
Lallemant, H.G.A. (1994), Pre-Cretaceous Tectonic Evolution of the Blue Mountains Province, Northeastern Oregon, USGS Professional Paper, 1438, 271-304.
Lambeck, K. (1988), Geophysical Geodesy, the Slow Deformations of the Earth, 718 pp., Clarendon Press, Oxford.
Leckie, D.A. (2001), Climatic controls on nonmarine sedimentation-maximum aggradation during lowstand, incision during transgession and highstand, in American Association of Petroleum Geology Annual Meeting, pp. 114, Denver.
Linacre, E., and B. Geerts (1997), Climates and Weather Explained, 432 pp., Routledge, London.
Little, T.A., S. Cox, J.K. Vry, and G.E. Batt (2005), Variations in exhumation level and uplift rate along the oblique-slip Alpine fault, central Southern Alps, New Zealand, Geological Society of America Bulletin, 117 (5/6), 707-723.
Luckman, B.H. (1976), Rockfalls and Rockfall Inventory Data: Some Observations from Surprise Valley, Jasper National Park, Canada, Earth Surface Processes, 1, 287-298.
MacKinnon, T.C. (1983), Origin of the Torlesse terrane and coeval rocks, South Island, New Zealand, Geological Society of America Bulletin, 94, 967-985.
Matsuoka, N. (1994), Diurnal freeze-thaw depth in rockwalls:field measurements and theoretical considerations, Earth Surface Processes and Landforms, 19, 423- 435.
Matsuoka, N. (2001), Direct observation of frost wedging in alpine bedrock, Earth Surface Processes and Landforms, 26, 601-614.
Matsuoka, N., and H. Sakai (1999), Rockfall activity from an alpine cliff during thawing periods, Geomorphology, 28, 309-328. 145
McArthur, J.L. (1987), The characteristics, classification and origin of late Pleistocene fan deposits in the Cass Basin, Canterbury, New Zealand, Sedimentology, 34, 459-471.
McCaffrey, R., M.D. Long, C. Goldfinger, P.C. Zwick, J.L. Nablek, C.K. Johnson, and C. Smith (2000), Rotation and plate locking at the southern Cascadia subduction zone, Geophysical Research Letters, 27 (19), 3117-3120.
McSaveney, M.J. (1992), A manual for weathering-rind dating of grey sandstones of the Torlesse Supergroup, New Zealand, 52 pp., Institute of Geological and Nuclear Sciences, Lower Hutt, New Zealand.
McSaveney, M.J. (2002), Recent rockfalls and rock avalanches in Mount Cook National Park, New Zealand, Reviews in Engineering Geology, 15, 35-70.
Miller, D.J., and T. Dunne (1996), Topographic perturbations of regional stresses and consequent bedrock fracturing, Journal of Geophysical Research, 101 (B11), 25523-25536.
Molnar, P., and P. England (1990), Late Cenozoic uplift of mountain ranges and global climate change: chicken or egg?, Nature, 346, 29-34.
Montgomery, D.R. (2002), Valley formation by fluvial and glacial erosion, Geology, 30 (11), 1047-1050.
Moon, B.P. (1984), Refinement of a Technique for Determining Rock Mass Strength for Geomorphological Purposes, Earth Surface Processes and Landforms, 9, 189-193.
Murton, J.B. (1996), Near-Surface Brecciation of Chalk, Isle of Thanet, South-East England: a Comparison with Ice-Rich Brecciated Bedrocks in Canada and Spitsbergen, Permafrost and Periglacial Processes, 7, 153-164.
Murton, J.B., J.-P. Coutard, J.-P. Lautridou, J.-C. Ozouf, D.A. Robinson, and R.B.G. Williams (2001), Physical Modelling of Bedrock Brecciation by Ice Segregation in Permafrost, Permafrost and Periglacial Processes, 12, 255- 266.
NIWA (2004), Database of daily temperature data for the central Southern Alps.
Norris, R.J., and A.F. Cooper (2000), Late Quaternary slip rates and slip partitioning on the Alpine Fault, New Zealand, Journal of Structural Geology, 23, 507- 520. 146
Orwin, J.F. (1998), The application and implications of rock weathering rind dating to a large rock avalanche, Cragieburen Range, New Zealand, New Zealand Journal of Geology and Geophysics, 41, 219-223.
Pack, R.T., and K. Boie (2002), Utah Rockfall Hazards Inventory: Phase I, 90 pp., Utah Department of Transportation, Salt Lake City, UT.
Pan, B., D.W. Burbank, Y. Wang, G. Wu, J. Li, and Q. Guan (2003), A 900 k.y. record of strath terrace formation during glacial-interglacial transition in northwest China, Geology, 31 (11), 957-960.
Pettinga, J.R., M.D. Yetton, R.J. Van Dissen, and G. Downes (2001), Earthquake source identification and characterisation for the Canterbury Region, South Island, New Zealand, Bulletin of the New Zealand Society for Earthquake Engineering, 34, 282-317.
Porter, S.C. (1975), Equilibrium-Line Altitudes of Late Quaternary Glaciers in the Southern Alps, New Zealand, Quaternary Research, 5, 27-47.
Rapp, A. (1960), Talus Slopes and Mountain Walls at Tempelfjorden, Spitsbergen A geomorphological study of the denudation of slopes in an arctic locality, Norsk Polarinstitutt Skr., 119 (4-7), 90-96.
Raymo, M.E., and W.F. Ruddiman (1992), Tectonic forcing of late Cenozoic climate, Nature, 359, 117-122.
Reidel, S.P., P.R. Hooper, G.D. Webster, and V.E. Camp (1992), Geologic Map of Southeastern Asotin County, Washington, Washington Division of Geology and Earth Resources.
Reidel, S.P., T.L. Tolan, P.R. Hooper, M.V. Beeson, K.R. Fecht, R.D. Bentley, and J.L. Anderson (1989), The Grande Ronde Basalt, Columbia River Basalt Group; Stratigraphic descriptions and correlations in Washington, Oregon, and Idaho, Geological Society of America Special Paper, 239, 21-54.
Rempel, A.W., J.S. Wettlaufer, and M.G. Worster (2004), Premelting dynamics in a continuum model of frost heave, Journal of Fluid Mechanics, 498, 227-244.
Reneau, S.L., and W.E. Dietrich (1991), Erosion Rates in the Southern Oregon Coast Range: Evidence for an Equilibrium between Hillslope Erosion and Sediment Yield, Earth Surface Processes and Landforms, 16, 307-322. 147
Rother, H., and J. Shulmeister (2005), Synoptic climate change as a driver of late Quaternary glaciations in the mid-latitudes of the Southern Hemisphere, Climate of the Past Discussions, 1, 231-253.
Saleeby, J.B., and Z. Foster (2004), Topographic response to mantle lithosphere removal in the southern Sierra Nevada region, California, Geology, 32 (3), 245-248.
Saltzer, R.L., and E.D. Humphreys (1997), Upper mantle P-wave structure of the eastern Snake River Plain and its relationship to geodynamic models of the region, Journal of Geophysical Research, 102, 11 829-11841.
Sass, O. (2005), Rock moisture measurements: techniques, results, and implications for weathering, Earth Surface Processes & Landforms, 30, 359-374.
Sass, O., and K. Wollny (2001), Investigations regarding alpine talus slopes using ground-penetrating radar (GPR) in the Bavarian Alps, Germany, Earth Surface Processes and Landforms, 26, 1071-1086.
Schmidt, K., and D.R. Montgomery (1995), Limits to Relief, Science, 270, 617-620.
Schutt, D.L., and C.E. Lesher (2005), The effect and seismic velocity of garnet and spinel lherzolite, Journal of Geophysical Research, in press.
Selby, M.J. (1980), A rock mass strength classification for geomorphic purposes: with tests from Antarctica and New Zealand, Zeitschrift fur Geomorphologie, 24 (1), 31-51.
Selby, M.J. (1982), Controls on the Stability and Inclinations of Hillslopes formed on Hard Rock, Earth Surface Processes and Landforms, 7, 449-467.
Selby, M.J. (1993), Hillslope Materials and Processes, 451 pp., Oxford University Press, Oxford.
Shulmeister, J., I. Goodwin, J. Renwick, K. Harle, L. Armand, M.S. McGlone, E. Cook, J. Dodson, P.P. Hesse, P. Mayewski, and M. Curran (2004), The Southern Hemisphere westerlies in the Australasian sector over the last glacial cycle: a synthesis, Quaternary International, 118-119, 23-53.
Smith, G.M., T.R.H. Davies, M.J. McSaveney, and D.H. Bell (2006), The Acheron rock avalanche, Canterbury, New Zealand-morphology and dynamics, Landslides, 3 (1), 62-72. 148
Speight, R. (1916), The Physiography of the Cass District, Transactions of the New Zealand Institute, 48, 145-153.
Statham, I. (1976), A Scree Slope Rockfall Model, Earth Surface Processes, 1, 43-62.
Statham, I., and S.C. Francis (1986), Influence of scree accumulation and weathering on the development of steep mountain slopes, in Hillslope Processes, pp. 245- 267, Allen & Unwin, Boston.
Stock, G.M., T.A. Ehlers, and K.A. Farley (2006), Where does sediment come from? Quantifying catchment erosion with detrital apatite (U-Th)/He thermochronometry, Geology, 34 (9), 725-728.
Suggate, R.P. (1990), Late Pliocene and Quaternary Glaciations of New Zealand, Quaternary Science Reviews, 9, 175-197.
Taber, S. (1929), Frost heaving, Journal of Geology, 37 (5), 428-461.
Thompson, S.M., and J. Adams (1979), Suspended load in some major rivers of New Zealand, in Physical Hydrology-New Zealand experience, edited by D.L. Murray, and P. Ackroyd, pp. 213-228, New Zealand Hydrological Society, Wellington.
Tippett, J.M., and P.J.J. Kamp (1993), Fission Track Analysis of the Late Cenozoic Vertical Kinematics of Continental Pacific Crust, South Island, New Zealand, Journal of Geophysical Research, 98 (B9), 16119-16148.
Tippett, J.M., and P.J.J. Kamp (1995), Geomorphic Evolution of the Southern Alps, New Zealand, Earth Surface Processes and Landforms, 20, 177-192.
Tolan, T.L., S.P. Reidel, M.V. Beeson, J.L. Anderson, K.R. Fecht, and F.J. Swanson (1989), Revisions to the estimates of the areal extent and volume of the Columbia River Basalt Group, Geological Society of America Special Paper, 239, 1-20.
Tucker, G.E., and R.L. Bras (1998), Hillslope processes, drainage density, and landscape morphology, Water Resources Research, 34 (10), 2751-2764.
Turcotte, D.L., and G. Schubert (2002), Geodynamics, 456 pp., Cambridge University Press, Cambridge.
Walcott, R.I. (1998), Modes of Oblique Compression: Late Cenozoic tectonics of the South Island of New Zealand, Reviews of Geophysics, 36 (1), 1-26. 149
Walder, J., and B. Hallet (1985), A theoretical model of the fracture of rock during freezing, Geological Society of America Bulletin, 96, 336-346.
Walker, G.W. (1979), Reconnaissance Geologic Map of the Oregon Part of the Grangeville Quadrangle, Baker, Union, Umatilla, and Wallowa Counties, Oregon, USGS Miscellaneous Investigations Series, Map I-1116.
Wandres, A.M., J.D. Bradshaw, S.D. Weaver, R. Maas, T.R. Ireland, and N. Eby (2004), Provenance analysis using conglomerate clast lithologies: A case study from the Pahau Terrane of New Zealand, Sedimentary Geology.
Wardle, P. (1991), Vegetation of New Zealand, 672 pp., Cambridge University Press, Cambridge.
Wellman, H.W. (1979), An uplift map for the South Island of New Zealand, and a model for the uplift of the Southern Alps, Bulletin of the Royal Society of New Zealand, 18, 13-20.
Wells, R.E., C.S. Weaver, and R.J. Blakely (1998), Fore-arc migration in Cascadia and its neotectonic significance, Geology, 26 (8), 759-762.
Wettlaufer, J.S., and M.G. Worster (1995), Dynamics of premelted films: Frost heave in a capillary, Physical Review E, 51 (5), 4679-4689.
Whipple, K.X., E. Kirby, and S.H. Brocklehurst (1999), Geomorphic limits to climate-induced increases in topographic relief, Nature, 401, 39-43.
Whipple, K.X., and B.J. Meade (2004), Controls on the strength of coupling among climate, erosion and deformation in two-sided, frictional orogenic wedges at steady state, Journal of Geophysical Research, 190 doi:10.1029/2003JF000019.
Whitehouse, I.E. (1981), A large rock avalanche in the Cragieburn Range, Canterbury, New Zealand Journal of Geology and Geophysics, 24, 415-421.
Whitehouse, I.E. (1983), Distribution of large rock avalance deposits in the central Southern Alps, New Zealand, New Zealand Journal of Geology and Geophysics, 26, 271-279.
Whitehouse, I.E. (1988), Geomorphology of the central Southern Alps, New Zealand: the interaction of plate collision and atmospheric circulation, Zeitschrift fur Geomorphologie Supplementeband, 69, 105-116. 150
Whitehouse, I.E., and G.A. Griffiths (1983), Frequency and hazard of large rock avalanches in the central Southern Alps, New Zealand, Geology, 11, 331-334.
Whitehouse, I.E., and M.J. McSaveney (1983), Diachronous talus surfaces in the Southern Alps, New Zealand, and their implications to talus accumulation, Arctic and Alpine Research, 15, 53-64.
Whitehouse, I.E., M.J. McSaveney, P.L.K. Knuepfer, and T.J.H. Chinn (1986), Growth of Weathering Rinds on Torlesse Sandstone, Southern Alps, New Zealand, in Rates of Chemical Weathering of Rocks and Minerals, edited by S.M. Colman, and D.P. Dethier, pp. 419-435, Academic Press, Orlando.
Wieczorek, G.F. (2002), Catastrophic rockfalls and rockslides in the Sierra Nevada, USA, Reviews in Engineering Geology, 15, 165-190.
Wilen, L.A., and J.G. Dash (1995), Frost Heave Dynamics at a Single Crystal Interface, Physical Review Letters, 74 (25), 5076-5079.
Willett, S.D. (1999), Orogeny and orography: The effects of erosion on the structure of mountain belts, Journal of Geophysical Research, 104 (B12), 28957- 28981.
Willett, S.D., and M.T. Brandon (2002), On steady states in mountain belts, Geology, 30 (2), 175-178.
Willett, S.D., R.L. Slingerland, and N. Hovius (2001), Uplift, shortening, and steady state topography in active mountain belts, American Journal of Science, 301, 455-485.
Worster, M.G., and J.S. Wettlaufer (1999), The Fluid Mechanics of Premelted Liquid Films, in Fluid Dynamics at Interfaces, edited by W. Shyy, pp. 339-351, Cambridge University Press, Cambridge.
Wunsch, C. (2004), Quantiative estimate of the Milankovitch-forced contribution to observed Quaternary climate change, Quaternary Science Reviews, 23, 1001- 1012.
Zhang, P., P. Molnar, and W.R. Downs (2001), Increased sedimentation rates and grain sizes 2-4 Myr ago due to the influence of climate change on erosion rates, Nature, 410, 891-897.
Zhu, C. (1996), Rates of Periglacial Processes in the Central Tianshan, China, Permafrost and Periglacial Processes, 7, 79-94. 151
Zhu, D.-M., O.E. Vilches, J.G. Dash, B. Sing, and J.S. Wettlaufer (2000), Frost Heave in Argon, Physical Review Letters, 85 (23), 4908-4911.