Climate and climate-related issues for the safety assessment SR-Site

Technical Report

TR-10-49

Climate and climate-related issues
for the safety assessment SR-Site
Svensk Kärnbränslehantering AB

December 2010

Svensk Kärnbränslehantering AB

Box 250, SE-101 24 Stockholm
Phone +46 8 459 84 00

TR-10-49

CM Gruppen AB, Bromma, 2010

Swedish Nuclear Fueland Waste Management Co

ISSN 1404-0344
Tänd ett lager:
SKB TR-10-49
P, R eller TR.

Climate and climate-related issues
for the safety assessment SR-Site
Svensk Kärnbränslehantering AB

December 2010

Keywords: Climate, Glacial, Ice sheet, Periglacial, Permafrost, Sea level,
Global warming, Safety assessment, SR-Site, SKBdoc 1263140.
A pdf version of this document can be downloaded from www.skb.se.

Preface
This document presents information on climate and climate-related issues relevant for long-term
safety of a KBS-3 repository for spent nuclear fuel. It supports the safety assessment SR-Site.
The report was compiled by Assoc. Prof. Jens-Ove Näslund, SKB. A number of additional authors
have contributed to various sections of the report as listed in Table 1-1.
The report was reviewed by Dr. Mike C. Thorne (Mike Thorne and Associates Limited), Prof. Peter
Jansson (Stockholm university) and Dr. Lillemor Claesson Liljedahl (SKB).

Stockholm, December 2010
Jens-Ove Näslund
Person in charge of the SKB climate programme

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3

Contents
1
1.1
1.2

1.3
1.4

1.5
1.6
2
2.1

2.2

2.3

2.4

3
3.1

3.2

Introduction
Purpose of the report
Strategy for managing climate-related issues in safety assessments
1.2.1 Climate and climate-related changes
1.2.2 Impact on repository performance and safety
1.2.3 Strategy for managing long-term evolution of climate-related
processes
Structure of the report
1.3.1 Overall report structure and authors
1.3.2 Structure for description of climate-related issues
Publications supporting the SR-Site Climate report
1.4.1 Publications produced for SR-Site
1.4.2 Publications produced for SR-Can, in parts or in whole used for
the SR-Site Climate report
Interactions between the climate programme and other SR-Site programmes
Summary of main changes made since the SR-Can Climate report

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The climate system
Components of the climate system
2.1.1 The atmosphere2.1.2 The hydrosphere2.1.3 The cryosphere2.1.4 The land surface
2.1.5 The biosphere
Climate forcing
2.2.1 Earth-orbital variations
2.2.2 Alteration of solar radiation
2.2.3 Tectonic processes
2.2.4 Anthropogenic forcing
Climate dynamics
2.3.1 Energy budget and radiation balance
2.3.2 The hydrological cycle
2.3.3 The carbon cycle
2.3.4 Feedback mechanisms and climate change
Climate in Sweden and Forsmark
2.4.1 Climate in Sweden
2.4.2 Climate at Forsmark

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Climate and climate-related issues
Ice sheet dynamics
3.1.1 Overview/general description
3.1.2 Controlling conditions and factors
3.1.3 Natural analogues/observations in nature
3.1.4 Model studies
3.1.5 Time perspective
3.1.6 Handling in the safety assessment SR-Site
3.1.7 Handling of uncertainties in SR-Site
3.1.8 Adequacy of references
Ice sheet hydrology
3.2.1 Overview/general description
3.2.2 Controlling conditions and factors
3.2.3 Natural analogues/observations in nature
3.2.4 Model studies
3.2.5 Time perspective
3.2.6 Handling in the safety assessment SR-Site

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3.3

3.4

3.5

4
4.1
4.2
4.3

4.4

4.5

5
5.1

6

3.2.7 Handling of uncertainties in SR-Site
3.2.8 Adequacy of references
Isostatic adjustment and shore-level changes
3.3.1 Overview/general description
3.3.2 Controlling conditions and factors
3.3.3 Natural analogues/observations in nature
3.3.4 Model studies
3.3.5 Time perspective
3.3.6 Handling in the safety assessment SR-Site
3.3.7 Handling of uncertainties in SR-Site
3.3.8 Adequacy of references
Permafrost
3.4.1 Overview/general description
3.4.2 Controlling conditions and factors
3.4.3 Natural analogues/observations in nature
3.4.4 Model studies
3.4.5 Time perspective
3.4.6 Handling in the safety assessment SR-Site
3.4.7 Handling of uncertainties in SR-Site
3.4.8 Adequacy of references
Surface denudation
3.5.1 Overview/general description
3.5.2 Controlling conditions and factors
3.5.3 Natural analogues/observations in nature
3.5.4 Model studies
3.5.5 Time perspective
3.5.6 Handling in the safety assessment SR-Site
3.5.7 Handling of uncertainties in SR-Site
3.5.8 Adequacy of references

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Climate and climate-related conditions for the SR-Site safety assessment
Rationale and general approach
Weichselian glacial history
4.2.1 Overview of Weichselian glacial history
4.2.2 Ice-marginal stillstands
Examples of Weichselian climates
4.3.1 Early Weichselian (117–74 kyrs BP)
4.3.2 Middle Weichselian (74–24 kyrs BP)
4.3.3 Late Weichselian including LGM (24–12 kyrs BP)
4.3.4 Climate variability during the last glacial cycle
Reconstructed last glacial cycle conditions at Forsmark
4.4.1 Ice sheet evolution
4.4.2 Shore-level evolution
4.4.3 Permafrost evolution
4.4.4 Combined reconstructed last glacial cycle parameters
Reference glacial cycle
4.5.1 Ice sheet evolution
4.5.2 Evolution of the Baltic Sea and the Forsmark shore-level
4.5.3 Permafrost evolution
4.5.4 Evolution of climate domains
4.5.5 Evolution of hydrological conditions and groundwater
4.5.6 Evolution of mechanical conditions
4.5.7 Surface denudation

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Additional climate cases for the safety assessment SR-Site
Global warming case
5.1.1 Background5.1.2 Ice sheet evolution
5.1.3 Shore-level evolution

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5.6

5.1.4 Permafrost evolution
5.1.5 Evolution of climate domains
5.1.6 Surface denudation
5.1.7 Exemplified climate conditions for the global warming case
Extended global warming case
5.2.1 Background5.2.2 Ice sheet evolution
5.2.3 Shore-level evolution
5.2.4 Permafrost evolution
5.2.5 Evolution of climate domains
5.2.6 Surface denudation
Extended ice sheet duration case
5.3.1 Background5.3.2 Ice sheet evolution
5.3.3 Surface denudation
Maximum ice sheet configuration case
5.4.1 Background5.4.2 Ice sheet evolution
5.4.3 Surface denudation
Severe permafrost case
5.5.1 Background5.5.2 Ice sheet evolution
5.5.3 Permafrost evolution
5.5.4 Surface denudation
Summary of climate cases for the SR-Site safety assessment

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List of abbreviations

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References

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5.2

5.3

5.4

5.5

Appendix 1

Air temperature data

301

Appendix 2

Ice sheet profiles

313

Appendix 3

Combination of buffer erosion and freezing

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1

Introduction

1.1

Purpose of the report

The purpose of this report is to document current scientific knowledge on climate and climate-related
conditions, relevant to the long-term safety of a KBS‑3 repository, to a level required for an adequate
treatment in the safety assessment SR-Site. The report also presents a number of dedicated studies on
climate and selected climate-related processes of relevance for the assessment of long term repository
safety. Based on this information, the report presents a number of possible future climate developments
for Forsmark, the site selected for building a repository for spent nuclear fuel in Sweden (Figure 1-1).
The presented climate developments are used as basis for the selection and analysis of SR-Site safetyassessment scenarios in the SR-Site main report /SKB 2011/. The present report is based on research
conducted and published by SKB as well as on research reported in the general scientific literature.

1.2

Strategy for managing climate-related issues insafety assessments

Swe

den

1.2.1 Climate and climate-related changes
Climate-related changes such as the development of permafrost and ice sheets and variations in shorelevel are the most important naturally occurring external factors affecting a repository for spent nuclear
fuel in a time perspective up to hundreds of thousands of years. Most of the processes occurring in
the biosphere and geosphere are likely to be affected by climate and climate-related changes in one
way or another.

Östhammar
municipality

Forsmark
Stockholm

Göteborg

Malmö

0

100

200

300

400 kmFigure 1-1. Location map.

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9

Over the past 2.5 million years, Scandinavia has experienced several cycles of glaciation e.g. /Ehlers
and Gibbard 2004/. Cold periods that include ice sheets gradually growing to maximum extents are
known as glacials. Intervening periods with warm climate, when the ice sheets have extents similar tothat at the present day, are called interglacials. A glacial cycle consists of a glacial and an interglacial.
Within the glacial periods, warmer interstadial periods occur, as well as cool stadial periods. The shifting
between glacials and interglacials is depicted for instance in the variation of the content of the heavy
oxygen isotope, 18O, in deep-sea sediments e.g. /Lisiecki and Raymo 2005/, which reflects both thetemperature of the sea and the volume of water that has been bound in land-based ice sheets andglaciers globally, see Figure 1-2.
Figure 1-2 shows that over the last 700 kyrs, glacial-interglacial cycles of about 100 kyr duration
have dominated global climate variation. The time frame of a Quaternary glacial-interglacial cycle,
around 100 kyrs, is similar to the time it takes for the radioactivity in the spent fuel to decay to levels
comparable to the activity in the natural uranium that was once used to manufacture the fuel /SKB
2011/. The time scale of a safety assessment for a final repository for spent nuclear fuel is one
million years. The 100 kyr time scale, corresponding to a first future glacial cycle, forms a basis forthe analysis of repository safety. For the remainder of the assessment period, this glacial cycle isassumed to be repeated. Alternative developments over the first 100 kyrs, including for instanceperiods of anthropogenic global warming, are also analysed. A time frame of 100 kyrs is also reasonable when determining the extremes of climate conditions the repository should be designed towithstand.
In very long time perspectives, several millions of years, geological processes like plate tectonic
movements and uplift or downwarping, additional to crustal displacements due to glacial loading
and unloading, will affect both the Earth climate system and the repository for spent nuclear fuel.
However, the SR-Site safety assessment address timescales up to one million years /SKB 2011/,
and therefore these geological processes are not included in the safety assessment.

1.2.2 Impact on repository performance and safety
Climate changes, permafrost development and the growth and decay of ice sheets will alter not only
surface but also subsurface conditions. Freezing of groundwater, the presence of ice sheets, and changes
in shore-level will change permeability, water turnover, groundwater pressure, groundwater flow
and chemical composition. The ice load will alter rock stresses during different phases of a glaciation.
This will, in turn, change bedrock permeability and as ice sheets melts away, a combination of large
horizontal stresses and high pore water pressures may cause glacially induced faulting. In general, theintegrated effects of the continuous evolution of climate-related conditions need to be considered, but

2

100 kyr cycles
2.5
3
3.5
4

5.5

5

4.5

4

3.5

3

2.5

Millions of Years ago

2

1.5

1

0.5

δ18O Benthic
carbonate (‰)

41 kyr cycles

4.5

0

Figure 1-2. Variations of δ18O in deep-sea sediments over the last 5.5 million years /Lisiecki and Raymo
2005/. High δ18O values indicate large global ice volume and low sea water temperature, i.e. colder climate
conditions, whereas low δ18O values indicate the opposite. The last 700 kyrs have been dominated by
100 kyr long glacial cycles. Modified from /Global Warming Art 2007, after Lisiecki and Raymo 2005/.

10

TR-10-49there are also a number of more specific phenomena of importance for repository safety that require
special attention. Based on the results of earlier and the present safety assessments, these include:•

The maximum hydrostatic pressure occurring at repository depth for glacial conditions.

•

The maximum permafrost and ground freezing depth.

•

The possible penetration of oxygen-rich groundwater to repository depth during glacial conditions.

•

The possible penetration of dilute groundwaters to repository depth during glacial and temperateconditions, potentially causing erosion of buffer and backfill.

•

The potential for glacially induced faulting.

•

Factors affecting retardation in the geosphere, like high groundwater fluxes and mechanicalinfluences on permeability.

1.2.3

Strategy for managing long-term evolution of climate-related processes

Definition of climate change and climate domains

There are various definitions of climate change. In the present report, climate change is definedaccording to the definition by the Intergovernmental Panel on Climate Change (IPCC), namely as“any change in climate over time whether due to natural variability or as a result of human activity”/IPCC 2001/. This differs from the definition used by United Nation’s Framework Convention on
Climate Change (UNFCCC) that defines climate change as “a change of climate that is attributed
directly or indirectly to human activity that alters the composition of the global atmosphere, and thatis in addition to natural climate variability over comparable time periods”. Since climate may change
due to both natural variability and to anthropogenic influences, the definition of climate change in
SR-Site follows the one of IPCC.
From a safety assessment point of view, it is not possible to predict a single climate evolution in a
100 kyr time perspective with enough confidence for a safety assessment, although some attempts
have been made to predict climate over this time scale, e.g. /Shaffer et al. 2009/. However, based on
knowledge of climate variations in the past and on inferred future climate change, the extremes within
which the climate of Sweden may vary can be estimated with reasonable confidence, also within
this long time frame. Within these limits, characteristic climate-related conditions of importance forrepository safety can be identified. The conceivable climate-related conditions can be represented asclimate-driven process domains /Boulton et al. 2001a/ where such a domain is defined as “a climatically determined environment in which a set of characteristic processes of importance for repository
safety appear”. In the following parts of the report these climate-driven process domains are referred
to as climate domains. The identified climate domains are denominated:• The temperate climate domain.
• The periglacial climate domain.
• The glacial climate domain.
The purpose of identifying climate domains is to create a framework for the assessment of issues ofimportance for repository safety associated with particular climatically determined environments that
may occur in Sweden. If a repository for spent nuclear fuel fulfils the safety requirements for plausible
climate developments, including the transitions between climate domains, then the uncertainty regarding the climate domains more precise extent in time and space is of less importance.
The temperate climate domain is defined as regions without permafrost or presence of ice sheets. It is
dominated by a temperate climate in a broad sense, with cold winters and either cold or warm summers.
Precipitation may fall at any time of the year. The precipitation falls either as rain or snow. The temperate
climate domain has the warmest climate of the three climate domains. Within the temperate climate
domain, a site may also at times be submerged by the sea. Climates dominated by global warming arealso included in the temperate climate domain.
The term periglacial is today used for a range of cold but non-glacial processes. In the periglacial
environment, permafrost is a central, but not defining, element /French 2007/. However, for thepresent work, the periglacial climate domain is defined strictly as regions that contain permafrost.

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Furthermore, the peri­glacial climate domain is a cold region but without the presence of an ice
sheet. In this climate domain, permafrost occurs either in sporadic (less than 50% spatial coverage),
discontinuous (between 50 and 90% coverage), or continuous form (more than 90% coverage).
Although true for most of the time, regions belonging to the periglacial climate domain are notnecessarily the same as regions with a climate that supports permafrost growth. For example, at theend of a period with periglacial climate domain the climate may be relatively warm, not building oreven supporting the presence of permafrost. Instead, permafrost may be diminishing. However, as
long as permafrost is present, the region is defined as belonging to the periglacial climate domain,
regardless of the prevailing temperature at the ground surface. This way of defining the climate
domain is used because the presence of the permafrost is more important for the safety function of
the repository than the actual temperature at the ground surface. In general, the periglacial climate
domain has a climate colder than the temperate climate domain and warmer than the glacial climate
domain. Precipitation may fall either as snow or rain. Within the periglacial climate domain, a site
may also at times be submerged by the sea.
The glacial climate domain is defined as regions that are covered by glaciers or ice sheets. Within
the glacial climate domain, the glacier or ice sheet may in some cases be underlain by sub-glacial
permafrost. In line with the definition of the periglacial climate domain, areas belonging to theglacial climate domain may not necessarily at all times have a climate that supports the presenceof ice sheets. However, in general, the glacial climate domain has the coldest climate of the three
climate domains. Precipitation normally falls as snow in this climate domain.
As previously mentioned, it is currently not possible to make confident long-term predictions of futureclimate, particularly taking into account the potential long-term significance of inferred current humaninduced perturbations of the natural climate system. It is, however, likely that all three climate domains
will appear repeatedly during the one million year assessment period, i.e. any reasonable evolution
will have to address them, and transitions between them.
Reference glacial cycle, climate cases and safety assessment scenarios

The climate domains are first used to describe a reference glacial cycle for the coming 120 kyrs
(Section 4.5), comprising a repetition of conditions reconstructed for the last glacial cycle (Section 4.4).
The reference glacial cycle is not to be seen as a prediction of a future climate development at
Forsmark, but as one example of a conceivable future evolution that covers climate-related conditions
and sequences that could be expected in a 100 kyr time perspective. Furthermore, using a reference
glacial cycle based on the last glacial cycle does not mean that glacial- and permafrost processes
are regarded as more probable than processes related to warm climates for e.g. the coming 100 kyrs.
The reference glacial cycle is supplemented by a number of additional climate cases that describe
alternative possible future developments of climate and climate-related processes at Forsmark(Figure 1-3). The selection and description of these climate cases are based on i) knowledge on pastchanges in climate and environmental parameters, ii) anticipated future climate change affected byanthropogenic action, and iii) knowledge as to which processes are of importance for repository safety.
A large part of the information used to select the climate cases has been obtained from previous safetyassessments. In SR-Site, six future climate cases are described, including the reference glacial cycle,
see Figure 1-3 and Section 4.5 and Chapter 5. For a summary of how the climate and climate-related
processes develop in the six climate cases, see Section 5.6. For a description of the rationale of using areference glacial cycle based on a reconstruction of conditions for the last glacial cycle, see Section 4.1.
The six climate cases are in turn used as the basis for the construction of safety assessment scenarios/SKB 2011/. In the safety assessment, the reference glacial cycle is used to construct a main scenario,
aiming at describing a reasonable evolution of the repository system given the timing and lengths ofperiods with different climate domains as derived from the reference glacial cycle. There are twovariants of the main scenario, i) a base case, comprising the reconstruction of climate and climaterelated processes as described in the reference glacial cycle, i.e. the repetition of reconstructed lastglacial cycle conditions, and ii) a global warming variant, defined from the global warming case,see Figure 1-3. The base case serves as one example of the evolution of climate, climate-related
processes and the repository for a glacial cycle dominated by natural climate change, i.e. a development not affected by anthropogenically cased global warming. The base case is complemented by the
global warming variant (Figure 1-3), where the inferred current human influence on climate is taken
into account.
12

TR-10-49

FUTURE

PAST

Reconstructions(present report)

Climate cases(present report)

Safety assessment
scenarios (TR-11-01)
SR-Site Main scenario

1. Reference glacial cycle(section 4.5)

Base case

2. Global warming(section 5.1)

Global warming variant
Selection of
additional
scenarios
based onsafetyfunctions

Reconstruction of lastglacial cycle conditions(section 4.3)

3. Extended global
warming (section 5.2)

4. Extended ice sheet
duration (section 5.3)

Additional
SR-Sitescenarios

5. Maximum ice sheet
configuration (section 5.4)

6. Severe permafrost(section 5.5)

Figure 1-3. Relationship between the reconstruction of last glacial cycle conditions, the reference glacial
cycle, the additional climate cases, and the corresponding SR-Site safety assessment scenarios. The redarrow indicates the choice of repeating the reconstructed last glacial cycle conditions for a future reference
glacial cycle. The black arrows indicate modifications made to the reference glacial cycle in order to construct additional future climate cases to obtain a comprehensive coverage of possible climate developmentsof relevance for long-term safety. Blue arrows show which climate cases have been used for the analysis ofwhich safety assessment scenario. The figure also shows in which sections the climate cases are describedin the present report. For information on the SR-Site safety assessment scenario selection (grey arrow) see
the text and /SKB 2011/.

Following the safety assessment analysis of the main scenario, a number of additional scenariosof importance for repository safety are selected (Figure 1-3). The scenario selection is based on theevolution in the main scenario and on an understanding of repository safety, see /SKB 2011/. The
latter is expressed through a number of safety functions and safety function indicators, used to make astructured selection of additional scenarios. Detailed information about the safety function indicatorsand how the scenario selections are made is found in /SKB 2011/. Following the scenario selection, theadditional scenarios are analyzed using the information from the associated climate cases (Figure 1-3).
The additional scenarios essentially address relevant situations not covered by the main scenario.
For example, for the glacial climate domain the maximum possible ice sheet thickness is one parameterof interest for the analysis of hydrostatic pressure at repository depth. In the base case of the main
scenario, the development of ice sheet thickness comes from a reconstruction of the Fennoscandian ice
sheet during the last glacial cycle (Section 3.1.4), whereas in an additional scenario the effect of ice
sheet thicknesses greater than those during the last glacial cycle are analysed using information from

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the maximum ice sheet configuration case (Section 5.4), see Figure 1-3. In this way, the main scenario
describes a plausible evolution of ice sheet thickness, and the additional scenario describes relevant
variations of this evolution.
It is important to include the climate domains and sequences that have the greatest impact on repository safety. It is also important to include sequences yielding the peak risk during the assessment
period. Based on results of earlier analyses, the highest risks (i.e. relating to radiogenic exposure tohumans) are likely to occur during periods of temperate climate. Typical situations are i) a terrestrial
system that has accumulated radionuclide releases over a long time, possibly at early stages in sea orlake sediments that subsequently is exposed and used for agriculture, and ii) a well intruding into thehost rock and used for domestic purposes. In addition, periods of periglacial climate domain needs
to be considered, where permafrost conditions change transport efficiency, exposure pathways andaccumulation processes, and also constrains the land-use and use of well water.

1.3

Structure of the report

1.3.1 Overall report structure and authors
Following the introduction chapter, the present report includes three main chapters:•

An overview description of the climate system (Chapter 2).

•

Identification and discussion of climate and climate-related issues, including reconstructions of
the identified variables for the last glacial cycle (Chapter 3).

•

A description of climate and climate-related conditions for the safety assessment reference glacial
cycle (Chapter 4).

•

A description of climate and climate-related conditions for the complementary climate cases(Chapter 5).

Chapter 2 includes an overview of present knowledge of the Earth’s climate system and the climate
conditions that can be expected to occur in Sweden on a 100 kyr time perspective. Based on this,climate-related issues relevant for the long-term safety of a KBS‑3 repository are identified. These
are documented in Chapter 3 to a level required for an adequate treatment in the safety assessment.
Subsequently, in Chapter 4, an evolution for a 120 kyr year period is presented, including discussionsof identified climate-related issues of importance for repository safety.
In Chapter 2 “The climate system” the intention is not to give a complete picture of the Earth climate
system but rather a summary sufficient for identification of climate-related conditions and processes
with a possible impact on repository performance. The interactions between these processes ofimportance for repository performance and other parts of the climate system set the limits for the
documentation of the climate system in this report.
In Chapter 3 there are five sections that, in detail, describe the climate-related phenomena within the
climate domains that may affect repository safety functions. These sections deal with the generalcharacteristics of ice sheets, glacial hydrology, shore-level displacement, permafrost, and surface
denudation, and for most of these parameters dedicated reconstructions of the development during
the last glacial cycle are presented.
In Chapter 4 the complete reconstruction of last glacial cycle conditions is presented, together withthe resulting reference glacial cycle over a future 120 kyr period. The latter thus includes the identified climate-related issues of importance for repository safety. Chapter 4 also includes a section with
examples of last glacial cycle climates, since the range and variability during the last glacial cycle
could be expected also for future glacial cycles.
In Chapter 5 the additional climate cases are described, complementing the reference glacial cycle
described in Chapter 4. For instance, cases of global warming are included here (including detailed
information for the Forsmark site), as well as other possible future climate developments.
In this context, a short comment on time scales should be made. Since Chapter 3 and Sections 4.1
to 4.4 mainly deals with reconstructions of climate-related conditions for the last glacial cycle thetime scales used here runs from 120 kyrs before present (BP) up to present day. In Section 4.5 and14

TR-10-49

onwards, on the other hand, the reconstructions from the previous parts are used to construct cases of
future developments of climate-related conditions. Therefore, the time scale used in Section 4.5 andonwards in the report runs from present day to 120 kyrs into the future.
In addition, the present report also includes three appendices. Appendix 1 and 2 describe data sets on
temperature and ice sheet profiles respectively, data that subsequently have been used within other
SR-Site activities. Appendix 1 gives a detailed account of the air- and ground temperature data
sets that have been produced for SR-Site, and have been used for e.g. ice sheet (Section 3.1.4) andpermafrost modelling /Hartikainen et al. 2010/. The appendix presents the original data, how the data
have been modified within SR-Site, and a discussion about data uncertainty. Appendix 2 describes indetail the 2D ice sheet profiles that are used for modelling of groundwater flow under glacial conditions /Vidstrand et al. 2010/ and for THM studies /Lönnqvist and Hökmark 2010/. The temperatureand ice sheet profile data sets are also presented in the /SKB 2010a/. Appendix 3 contains studies ona hypothetical situation where freezing at repository depth and bentonite buffer erosion are combined,
complementing the information in the buffer freezing scenario in the SR-Site Main report.
As further described in the SR-Site Main report and in /SKB 2010b/ the content of the present
report has been audited by comparison with national FEP databases from other assessment projects
as compiled by NEA (Nuclear Energy Agency), see the SR-Site FEP-report. The models used forthe construction of the reference glacial cycle and some of the complementary climate cases are
documented in the SR-Site Model summary report /SKB 2010g/, while the data produced by themodels are presented in the SR-Site Data report /SKB 2010a/. Results from the present climate
report are used in the SR-Site Main report /SKB 2011/ and in reports supporting the main report.
The present report follows as far as possible the SR-Site template for documentation of processes
regarded as internal to the repository system. However, the term processes is typically not used inthis report, instead the term issue has been used. Each issue includes a set of processes together
resulting in the behaviour of a system or feature. For instance, “ice sheet dynamics” is the result of
several climatological, hydrological, and mechanical processes. But as the ice sheet can be seen as
one entity in its interaction with the bed, these processes are jointly accounted for under the heading“ice sheet dynamics”. Given the content of the report, and the aim of following the template for
SR-Site process reports, there is some necessary repetition of text between some of the chaptersand sections.
A list of abbreviations is found at the end of the report.
The experts involved in describing the basic information about the climate system and climate-related
issues are listed in Table 1-1. In addition, the sections “Handling in the safety assessment SR-Site” and“Handling of uncertainties in SR-Site” have been produced by Jens-Ove Näslund, SKB, in collaboration with the expert that assembled the information on the issue in question. All experts contributing
to the present report are included in the SR-Site expert database. The sections The climate system,
Climate variability and change, Climate in Sweden and Isostatic adjustment and shore-level displacement were originally written for the SR-Can safety assessment /SKB 2006a/ by the authors listed inTable 1-1, and have been adapted for the present report by Jens-Ove Näslund, SKB.

1.3.2 Structure for description of climate-related issues
In discussing each of the various climate-related issues, the following standardised structure has
been adopted.
Overview/general description

Under this heading, a general description of the current knowledge regarding the climate-related
issue is given.
Controlling conditions and factors

The external and internal conditions and factors that control each “issue” are discussed. External refers
to systems that are not part of the described system/feature and that the described system/feature interacts
with. For example, for “ice sheet dynamics”, the atmosphere and the bed are external factors andrelevant aspects of them are described. Internal refers to conditions and factors governing system/feature
TR-10-49

15

Table 1-1. Contributors to individual chapters and sections in the present report.
Chapter

Section(s)Subsection(s)Author(s)1. Introduction
2. The climate
system

1.1–1.6
2.1 Components ofthe climate system

All
All, except 2.1.3

Jens-Ove Näslund, SKB
Jörgen Bogren and Torbjörn Gustavsson, Klimator AB,
Gothenburg University
Jörgen Bogren and Torbjörn Gustavsson, Klimator AB,
Gothenburg University
Jens-Ove Näslund, SKB
Jörgen Bogren and Torbjörn Gustavsson, Klimator AB,
Gothenburg University
Jens-Ove Näslund, SKB
Jörgen Bogren and Torbjörn Gustavsson, Klimator AB,
Gothenburg University
Moberg et al. 2006, Stockholm university
Jens-Ove Näslund, SKB
Hans Alexandersson, Swedish Meteorological andHydrological Inst. (SMHI)

2.1.3

2.2 Climate forcing

2.3 Climate
dynamics

2.2.1, 2.2.3
2.2.2, 2.2.4
All, except 2.3.4
2.3.4

2.4 Climate in
Sweden andForsmark
3. Climate andclimate-related
issues3.

3.

3.1 Ice sheet
dynamics
3.2 Ice sheethydrology

3.3 Isostatic
adjustment andshore-levelchanges

3.4 Permafrost

2.4.1
2.4.2
All

Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB

All, except 3.2.4,
3.2.6, 3.2.8
3.2.4

Peter Jansson, Stockholm university

3.2.6, 3.2.8
3.3.1, 3.3.7

3.3.2, 3.3.4, 3.3.5
3.3.3, 3.3.6, 3.3.8
3.4.1, 3.4.4, 3.4.5,
3.4.7
3.4.2, 3.4.3

3.5 Surface
denudation
4. Climate and
4.1 Rationale andclimate-related
general approachconditions for the
SR-Site safetyassessment
4.2 Weichselian
glacial history

4.3 Examplesof Weichselian
climates

3.4.6, 3.4.8
3.5.1–3.5.4,
3.5.5–3.5.8
All

4.2.1

4.2.2
All, except 4.3.2
section on EarlyMIS 3, 4.3.2
section on Middle
MIS 3
4.3.2 Early MIS 3
4.3.2 section on
Middle MIS 3 and4.3.3. section on
Late Weichselian

16Peter Jansson, Stockholm university
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB
Pippa Whitehouse, University of Durham
Pippa Whitehouse, University of DurhamJens-Ove Näslund, SKB
Juha Hartikainen, Aalto University School of Science andTechnology
Jens-Ove Näslund, SKB
Juha Hartikainen, Aalto University School of Science andTechnology
Jens-Ove Näslund, SKB
Mats Olvmo, Gothenburg university
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB

Jens-Ove Näslund, SKB
Hanna Lokrantz, Bergab*
Gustav Sohlenius, SGU*
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB

Karin Helmens, Stockholm University**
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB
Erik Kjellström, Rossby Centre, Swedish Meteorologicaland Hydrological Inst. (SMHI)***
Jenny Brandefelt, KTH***
Gustav Strandberg, Rossby Centre, Swedish
Meteorological and Hydrological Inst. (SMHI)***

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Chapter

5. Additional
climate cases
for the safetyassessment
SR-SiteSection(s)Subsection(s)Author(s)4.4 Reconstructedlast glacial cycle
conditions at
Forsmark
4.5 Reference
glacial cycle
5.1 Global warming
case

All

Jens-Ove Näslund, SKB

All

Jens-Ove Näslund, SKB

All, except 5.1.7

Jens-Ove Näslund, SKB

5.1.7

Jens-Ove Näslund, SKB
Erik Kjellström, Rossby Centre, Swedish Meteorologicaland Hydrological Inst. (SMHI)***
Jenny Brandefelt, KTH Royal Institute of Technology***
Gustav Strandberg, Rossby Centre, Swedish
Meteorological and Hydrological Inst. (SMHI)***
Jens-Ove Näslund, SKB

5.2 Extended global All, except
warming case
5.2.3 section on
Sea-level rise up
to year 2100
5.2.3 section on
Sea-level rise up
to year 2100
5.3 Extended ice
Allsheet duration case
5.4 Maximum ice
Allsheet configurationcase
5.5 Severe
All, except 5.5.3
permafrost caseAppendix 1 – Air
temperature data

5.6 Summary ofclimate cases for
the SR-Site safetyassessment
AllAppendix 2 – Ice All
sheet profiles
Appendix 3 –Combination of
buffer erosion
and freezing

Jens-Ove Näslund, SKB
Lars Brydsten, Umeå University****Anders Engqvist, KTH Royal Institute of Technology****Tobias Lindborg, SKB****
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB

5.5.3

Juha Hartikainen, Aalto University School of Science andTechnology*****
Jens-Ove Näslund, SKB
Jens-Ove Näslund, SKB

All

Jens‑Ove Näslund, SKB

All

Jens‑Ove Näslund, SKB

A3.1

Jens-Ove Näslund, SKB

A3.2

Juha Hartikainen, Aalto University School of Science andTechnology******
Martin Birgersson, ClayTech
Jens-Ove Näslund, SKB

A3.3
A3.4

Authors with asterisks have not directly contributed written parts to the report. Instead texts have been adapted byJens-Ove Näslund from the following supporting reports:* /Lokrantz and Sohlenius 2006/.
** /Helmens 2009a/.
*** /Kjellström et al. 2009b/.
**** /Brydsten et al. 2009/.
***** /Hartikainen et al. 2010/.
****** /SKB 2006a, Section 4.4.1/.

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17

behaviour and that are generally included in models of the system/feature, e.g. for “ice sheet dynamics”
the creep process of ice. The external and internal conditions and factors are those that relate to thedescribed behaviour of the system/feature, e.g. for “ice sheet dynamics” air temperature, geothermalheat flow and ice properties.
A table summarising the influence of the climate-related issue on the geological barrier of the repositoryis presented. The table includes the geosphere variables influenced by the system/feature and the variables
associated with the system/feature that are required to be known to quantify the interaction between thesystem/feature and the repository system, e.g. for “ice sheet dynamics” basal temperature and melt-rate.
In this context, the repository system includes both the geosphere and the engineered barriers.
Natural analogues/observations in nature

Under this heading observations in nature and, when applicable, present-day natural analogues
regarding the process are summarised.
Model studies

In this section model studies of the process are summarised. This documentation focuses on process
understanding and for instance includes sensitivity analyses. The handling of external and internal
controlling conditions and factors are described, e.g. if spatially and temporally varying internal
conditions considered.
Time perspective

The time scales on which the system/feature operates and changes are documented.
Handling in the safety assessment SR-Site

The handling of the documented interactions with the repository is discussed. As a result of theinformation under this subheading, a mapping of the climate-related issues to method of treatment and,in relevant cases, applicable models are produced. The mapping is characterised on various time scales.
Handling of uncertainties in SR-Site

Given the adopted handling in the safety assessment SR-Site as described above, the treatment ofdifferent types of uncertainties associated with the issue, within that general framework, is summarised
under the following headings.
Uncertainties in mechanistic understanding: The uncertainties in the general understanding of theissue are discussed based on the available documentation and with the aim of answering the question:
Are the basic scientific mechanisms governing the issue understood to a level necessary for thesuggested handling? Alternative models may sometimes be used to illustrate this type of uncertainty.
Model simplification uncertainties: In most cases, the quantitative representation of a process containsimplifications, for instance in a numerical model. These may result in a significant source of uncertainty
in the description of the system evolution. Alternative models or alternative approaches to simplification for a particular conceptual model may sometimes be used to illustrate this type of uncertainty.
Input data and data uncertainties: The input data necessary to quantify the process for the suggested
handling are documented. The treatment of important input and output data and associated data
uncertainties are described in detail in /SKB 2010a/, to which reference is made as appropriate.
Adequacy of references

For the description of each climate-related issue, the qualification of key references are described.

18

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1.4

Publications supporting the SR-Site Climate report

1.4.1 Publications produced for SR-Site
Within the SKB climate programme, about 15 supporting reports and publications have been produced
to support the SR-Site Climate report. Many of the topics that have been studied are by necessity linked
to each other and in some cases therefore partly overlapping. Studies on climate, palaeoclimate andWeichselian glacial history were conducted by /Helmens 2009a, Houmark-Nielsen 2009, Kjellström
et al. 2009b (including erratum Feb 2010), Wohlfarth 2009/. A special issue of Boreas /Wohlfarth andNäslund 2010/, with papers on conditions in Scandinavia during a long period prior to the Last Glacial
Maximum of the last glacial cycle, was published as a result of a workshop on this topic arranged by
SKB in 2007 /Näslund et al. 2008/. Permafrost development was studied by /Hartikainen et al. 2010/
while glacial hydrology was studied by /Jansson and Näslund 2009/. Surface denudation (i.e. weathering and erosion) was investigated by /Olvmo 2010/, while bore hole temperatures, geothermal heatand palaeoclimate was studied by /Sundberg et al. 2009/. Various aspects of glacial isostatic adjustment
have been studied, such as sea-level change and shore-level displacement /Whitehouse 2009/ andcrustal stresses under glacial conditions /Lund et al. 2009, Lund and Näslund 2009/. In addition,/Brydsten et al. 2009/ studied extreme sea-levels at Forsmark up until year 2100.
The climate studies performed for SR-Site have resulted in a number of scientific publications, orcontribution to such publications, published in peer-reviewed journals, see /Whitehouse et al. 2006,
Helmens et al. 2007a, Engels et al. 2008a, 2008b, Bos et al. 2009, Brandefelt and Otto-Bliesner 2009,
Clark et al. 2009, Helmens et al. 2009b, Väliranta et al. 2009, Engels et al. 2010, Helmens and Engels
2010, Kjellström et al. 2010a, Ampel et al. 2010, Wohlfarth 2010, Strandberg et al. 2010, Brandefelt
et al. 2011, Van Meerbeeck et al. submitted/.
Furthermore, the following conference abstracts have been presented from the SR-Site climate
studies /Engels et al. 2006, Bos et al. 2007, Brandefelt and Strandberg 2007, Helmens et al. 2007b,
Helmens et al. 2007c, Kaislahti-Tillman et al. 2007, Väliranta and Helmens 2007, Engels et al.
2008c, 2008d, Helmens et al. 2008, Kjellström et al. 2008, Strandberg et al. 2008, Väliranta et al.
2008, Helmens 2009b, Helmens et al. 2009a, Kjellström et al. 2009a, Strandberg et al. 2009a, b,
Brandefelt et al. 2010, Helmens et al. submitted, Shala et al. submitted/.
In addition to the above SR-Site studies, a large number of papers from the general scientific literature
have been used as input to the present SR-Site Climate report.

1.4.2

Publications produced for SR-Can, in parts or in whole used for the
SR-Site Climate report
In addition to the reports and publications produced for SR-Site described above, some of the publications produced for the SR-Can safety assessment have in parts or in whole been used also for SR-Site.
This includes parts of the SR-Can Climate report /SKB 2006a/, and studies on climate /Hohl 2005,
Moberg et al. 2006/, Weichselian glacial history /Lokrantz and Sohlenius 2006/, glacial hydrology/Jansson et al. 2007/, glacial erosion /Påsse 2004/, and shore-level displacement /Påsse 2001/.

1.5Interactions between the climate programme and other
SR-Site programmes

Data and information from the climate programme has been used within many other parts of the SR-Sitesafety assessment. The main flow of data within the climate programme and to other SR-Site activities
are summarised in Figure 1-4 and below.
Data on air temperature have been used to simulate the development of ice sheets (Section 3.1.4) andpermafrost, see Section 3.4.4, 5.5.3 and /Hartikainen et al. 2010/ and for evaluation of thermal
gradients and geothermal heat flow /Sundberg et al. 2009/. Data on changing ice sheet configurations
over time have in turn been used for studies of variations in stress in the Earth crust /Lund et al.
2009/, which are used for assessing the potential for and consequences of glacially induced faulting/SKB 2010c/. Ice sheet data are also needed for simulations of groundwater flow under glacial conditions /Vidstrand et al. 2010/. Data from permafrost simulations have been used to make appropriate
TR-10-49

19

Shore-level
displacement

Landscape modelling(e.g. lake formation)
Crustal stress modelling(glacially induced faulting)

Ice thickness andsurface gradients

THM processes, e.g.
calculation of hydraulicjackling

ClimateMaximum permafrost
depth
Bedrock temperatures
Ground surface
temperature

Ground- and surface waterflow modelling for glacialand permafrost conditions

Geothermal heat flow
Thermal gradients

SR-Site scenario on
repository freezing

Figure 1-4. Main flow of climate- and climate-related data in SR-Site. Coloured boxes show data within
the climate programme. White boxes show other SR-Site activities, using data from the climate programme.
THM processes=Thermo-hydro-mechanical processes.

selections of conditions when studying groundwater flow under permafrost conditions /Vidstrand
et al. 2010/ as well as for studies of surface hydrology under a colder climate /Bosson et al. 2010,
SKB 2010d/. Data on shore-level change have been used for a description of the past and future
evolution of the landscape as influenced by e.g. the isostatic recovery from the last glacial cycle/Söderbäck 2008, SKB 2010d, e/, for instance in order to model future lake formation. Shore-level
data are also used for studying groundwater flow in potential future situations when the isostatic uplift
has progressed further. Data on permafrost and glacial development have also been used for investigating THM processes /Lönnqvist and Hökmark 2010/ and for assessing the potential for freezing ofvarious parts of the repository /SKB 2011/. Climate data were also included in the site selection
process /SKB 2010f/.

1.6

Summary of main changes made since the SR-Can
Climate report

Below is a summary of main changes in the SR-Site Climate report compared to the SR-Can Climate
report /SKB 2006a/. In addition to these major changes, all chapters and sections of the report havebeen updated.
•

Complementary and up to date studies on Weichselian glacial history are included, used for the
motivation and description of the reconstruction of the last glacial cycle and the SR-Site reference
glacial cycle.

•

In order to better describe possible future climate variability and to exemplify climates with thedifferent climate domains, dedicated climate model simulations have been performed for selectedperiods during the last glacial cycle and for a future climate dominated by global warming.

•

Detailed site-specific permafrost simulations have been performed in 2D, with an improved
permafrost model, complementary to the 1D permafrost simulation performed for SR-Can.

20

TR-10-49•Complementary information on shore-level displacement from 3D GIA simulations is included.

•

Uncertainties in the palaeotemperature curves used for e.g. ice sheet- and permafrost simulations
are better described.

•

A major update in the description of ice sheet hydrology have been made in accordance withrecent literature.

•

Knowledge and descriptions of possible future climate changes have been updated according torecent literature.

•

The number of possible future climate evolutions for further analysis in the safety assessment
has been increased from four to six.

•

A dedicated description of transitions between climate domains have been included.

•

Detailed information and evaluation of the temperature proxy data used for simulations of ice
sheets and permafrost have been included (Appendix 1).

•

Detailed information on ice sheet profiles used in geohydrological studies have been included(Appendix 2).

•

In SR-Can, the process of surface denudation (i.e. surface erosion and weathering) was reported
in the Geosphere process report /SKB 2006b/. In SR-Site, this process has been moved to the
Climate report (Section 3.5). In addition, this section now includes a study on denudation for aregion centred at Forsmark.

•In SR-Site, crustal stress studies for analysis of glacially induced faulting are not included in the
Climate report as was done in SR-Can. Instead, these studies are reported in /SKB 2010c/.

•

A new complementary study on the effect of freezing of water in bentonite buffer clay cavities
have been included.

•

Because the SR-Site safety assessment only concerns the Forsmark site, the present report describes
climate and climate-related issues only for the Forsmark site. The SR-Can Climate report
described both the Forsmark and the Laxemar site.

TR-10-49

21

2

The climate system

Weather and climate has great impact on Earth’s life and environment. The always changing weather
is determined by the fluctuating state of the atmosphere surrounding us, and characterised anddescribed by weather elements such as air temperature, air pressure, humidity, clouds, precipitation,
visibility and wind. The predictability of weather is limited. The reason for this is that the atmosphericcharacteristics and processes determining the weather can only be predicted over periods from a few
hours up to about two weeks.
The climate is by definition a summary of weather conditions during a certain period of time within acertain area. Climate is usually described as mean values of the weather parameters and their variability
over a specific time-span and covering a distinct area. The climate varies between different places due
to factors such as latitude, altitude, distance to the ocean and vegetation. The climate also varies overtime on different time scales from decadal, through centennial, and millennial to several thousands andmillions of years. The complex range of climatic variability is a result of external forcing acting on
the Earth’s climate system as well as of internal factors acting within the system. The climate system
includes five major components: the atmosphere, the hydrosphere, the cryosphere, the land surface andthe biosphere (Figure 2-1). Within the climate system, there are many physical, chemical, and biological processes that interact on a wide range of temporal and spatial scales both within and between
the components.
The general descriptions of the climate system in Chapter 2 is based on information from /Ahrens
1994, Bogren et al. 1998, Burroughs 2001, Ruddiman 2001, Weart 2003/. More specific information
is based also on other references as mentioned in the text.

Changesin Solar
input

Changes inatmospherecompositionand circulation

Changes inhydrologicalcycle
CloudsPrecipitation

AlbedoAtmosphere

Evaporation

Ice sheets,
Glaciers,
Snow coverHeat
exchange

Sea iceOcean
coupling

H2O N2 O2 CO2 O2 etc
Aerosols

Terrestrial
radiation

Human
influences

Ocean

Changes in ocean
circulation andbiogeochemistryAir-land-biomass
coupling

Land
Changesin land use,
vegetation,
ecosystems

Figure 2-1. Schematic view of the components of the global climate system, their processes, interactions
and some aspects that may change (boxes). Modified from /NSF 2000/.
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23

2.1

Components of the climate system

2.1.1

The atmosphere

The aim of this section is to give a brief overview of the main components of the climate system, inorder to set the context for readers not familiar with climatology, ice sheets, permafrost etc, and also
to set the rest of the report content in a wider framework.

Overview/general characteristics

The atmosphere is the mixture of gases that surrounds the Earth, being held there by gravity. Abouthalf of its mass is to be found in the lowest 5–6 km and more than 99% of the mass is within 40 km of
the surface of the Earth. This can be compared with the radius of the Earth which is about 6,400 km.
The atmosphere is generally considered to be c. 1,000 km thick. However, it has no distinct outer border
but a successive decrease of density as it merges into the inter-planetary medium. The origin of theatmosphere is gases that were released from the interior of the Earth by, for example, volcanic eruptions. The atmosphere has undergone major changes during the evolution of the Earth due to several
factors, in which changes of solar radiation and biological activities have played important roles.
Oxygen was not added to the atmosphere until produced by living organisms through photosynthesis,
an event that began in the late Archaean, ~2.7 billion years ago.
The atmosphere can be divided into four major layers: the troposphere, the stratosphere, the mesosphereand the thermosphere (Figure 2-2). This division is based upon the temperature structure of the atmosphere. In the lowest layer, the troposphere, all processes that we know as weather take place. The
troposphere contains almost all of the water vapour and clouds, and approximately 75% of the total
atmospheric mass. The troposphere is characterised by a more or less uniform decrease in temperature
with height, averaging 0.65°C per 100 m. Generally, it has an increasing wind speed with height
except close to the ground surface where topography affects the wind. The upper part of the troposphere is limited by the tropopause, which is very important since it prevents convective, vertical air
movement. The distance from the Earth surface to the tropopause varies according to latitude andseason ranging from 16–17 km at the poles to 8–9 km at the equator.

100

Thermosphere

90

Mesopause
80

70

Altitude (km)Mesosphere
60

50Stratopause
4030Stratosphere20

10Tropopause
Troposphere

0

-100

-80

-60

-40

-20

0

20

Temperature (°C)

Figure 2-2. Vertical structure of the atmosphere.

24

TR-10-49

Composition

The natural atmosphere can be divided into three types of constituents: permanent gases, variablegases, and non-gaseous constituents. Of the permanent gases, the main components are nitrogen(N2), 78.1% volume mixing ratio, oxygen (O2), 20.9% volume mixing ratio and argon (Ar) 0.93%
volume mixing ratio. These gases are mostly passive in meteorological processes, they have only
limited interactions with the incoming solar radiation and they do not interact with the infrared radiation emitted by the Earth. However, oxygen is highly active chemically; it can combine with othersubstances and also form ozone (O3).
The variable gases on the other hand, such as carbon dioxide, methane, nitrous oxide and also ozone,
absorb and emit infrared radiation. They are the most important components in terms of meteorological processes. These gases are also called greenhouse gases, and they are sparsely occurring with a
total volume mixing ratio in dry air of less than 0.1% by volume. However, these gases play anessential role in the energy budget of the Earth/atmosphere system (Figure 2-3). The atmosphere also
contains the greenhouse gas water vapour with a volume mixing ratio that is highly variable having atypical value of about 1%. These greenhouse gases absorb the infrared radiation emitted by the Earth.
Thus they tend to raise the temperature close to the Earth’s surface. The greenhouse gases water
vapour, carbon dioxide, and ozone also absorb solar short-wave radiation. Water vapour is the strongest
greenhouse gas and has a fundamental effect on climate as it exhibits strong spatial and temporalvariations in concentration.
Reflected
solar radiation

Incoming
solar radiation

Emitted infrared
radiation

107

342

77

235

Temperature at 5 km -19oC

67

195

40

Emission from andabsorption in the atmosphere

Absorption in
the atmosphereSensible Latent
heat
heat
30
168

Temperature at
Earth surface +15oC

324

390

24

78

Figure 2-3. Earth’s radiation and energy balance. Units are expressed in Wm–2.

TR-10-49

25

In the atmosphere, there are also non-gaseous constituents including dust, smoke and salt particles
often referred to as aerosols (solid and liquid particles). These are present in different concentrations
from a few million particles per cubic metre of air and they can affect the transmission of radiation
through the atmosphere. The effect depends on particle characteristics such as shape, size and colour,
and also on the particle distribution and their altitude above the surface. Sulphate aerosols are ofgreat importance in the atmosphere, since they directly affect cloud formation, and also on the sizeof the cloud droplets formed, which in turn has an influence on cloud albedo. The content of aerosolsin the atmosphere is altered by such things as volcanic eruptions and human activities. Althoughmost climate-related effects of particles arise below the tropopause, volcanic eruptions can ejectparticles into the stratosphere.
Dynamics – atmospheric circulation

The atmosphere is in continuous motion. The underlying cause of the atmospheric circulation is thepressure differences caused by the unequal distribution of solar radiation over the Earth’s surface andthe unequal heating of land and ocean areas. Averaged over the entire Earth, incoming solar radiation
equals outgoing radiation. This energy balance is not maintained for each latitude, since the tropics
experience a net radiative gain in energy whereas the Polar Regions suffer a net loss. These inequities
are balanced when heat is transported from equatorial regions pole-ward.
The thermal energy in the atmosphere is transferred in different ways; by convection, conduction,
latent heat transfer, advection and radiation. Convection means that the heated atmosphere is physicallytransported from one location to another in, for example, rising thermals. By conduction, heat energy
is transferred through the medium by molecular impact, whereas the medium itself does not move.
Latent heat transfer means that the heat energy absorbed by the melting or evaporation of water in one
location may be released elsewhere in the atmosphere when the water vapour subsequently condenses
or the water freezes. Heat energy may be transferred horizontally by advection when the winds cause
the mass movement of the atmosphere. Finally, energy is transferred in the atmosphere through theemission and absorption of radiation, moving through the space at the velocity of light.
The motion of the atmosphere is determined by two types of forces; driving forces and steering forces
resulting in the pattern of winds. The driving forces exist regardless of whether or not the air is moving.
Vertically, gravity is acting downwards whereas the vertical pressure gradient, with lower pressures athigher altitude, tend to force air upwards. Gravity and the vertical pressure gradient are generally close
to being in balance. Horizontal pressure gradients forcing air movement are also present, and are both
spatially and temporally variable on a range of scales. The steering forces are related to the motionof the air. The Coriolis effect, due to the rotation of the Earth, deflects the air from its initial motion
generated by the driving forces. Frictional forces depend on the degree of roughness of the surface theair moves over and act to reduce the wind speed resulting from the prevailing pressure gradient and to
produce a component of flow across the isobars. Around curved isobars, the centripetal acceleration
produces a force that deflects the inward motion.
The general circulation of the atmosphere is seen in Figure 2-4. The circulation of the troposphereconforms to a pattern that is characterised by low pressure areas in the tropics where warm air rises.
The rising air spreads towards the poles in the upper troposphere and eventually descends in the
middle latitudes. This forms part of the so-called Hadley cells. The sinking air leads to the development of large areas dominated by subtropical high pressures. These high-pressure cells are stableand occur over oceans and continents. From these high-pressure areas the trade winds have their
origin. Over the poles the coldest air masses, with dense relatively stable air, are found. The cold anddense polar air drifts towards the middle latitudes where, in contact with the warmer air, it gives rise
to cyclonic activity. The interaction of cold and warm air masses in the middle latitudes is largely
influenced by a belt of high-speed winds at high altitude in the upper troposphere.
The semi-stationary polar front (Figure 2-4) is where cold polar air masses and warm tropical air
masses meet. When these air masses move, large horizontal undulations, Rossby waves, form in
the polar front. Due to the temperature gradients and moving air masses, the polar front generates
mid-latitude low-pressure cyclone systems. Further north, where the polar air mass meets the evencolder arctic air mass, a similar arctic front is formed, which may be as strong as the polar front.

26

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Polar highPolar easterlies

Polar front
Subtropical highs

Westerlies

Mid-latitude

NE tradewinds

Hadley cell
Intertropicalconvergence
zone

0°
SE
trade winds
30°
Westerlies

60°

Figure 2-4. The general circulation of the atmosphere, shown as a three-cell circulation model. The atmo­
sphere is the most unstable and rapidly changing part of the climate system.

Semi-permanent centres of either high or low air pressure, associated with the polar front, characterise
particular geographical areas. For instance, during the winter season, the Icelandic Low is located fromnear Iceland and stretches into the Barents Sea. It is associated with frequent cyclone activity, bringing
precipitation in over Fennoscandia. In spring, when the arctic pressure gradients decrease, the Icelandic
Low and other similar centres of low pressure weaken. In line with this, the so-called Siberian High,a prominent high pressure cell located over Eurasia in winter, is also weakened and replaced by a low
pressure system during summer. In summer, pressure gradients are typically weaker than during winter,with less cyclonic activity as result. These semi-permanent centres of high or low pressure, and theirseasonal variation, affect the entire weather pattern in the higher latitudes of the Northern Hemisphere,
and in turn also the hydrosphere and cryosphere in these regions.

2.1.2

The hydrosphere

General structure

The hydrosphere comprises all liquid water that can be found on Earth. It includes groundwater, surface
water in fresh-water lakes and watercourses and saline water in the oceans. Approximately 70% of the
Earth’s surface is covered by oceans, which make the oceans a very important part of the hydrosphere,
and they play a major role within the climate system. The oceans have a large capacity for storing andtransporting heat energy, and also serve as storehouses for carbon dioxide. Due to the large thermal
inertia of the oceans, they play the role of a dampener of temperature changes.
One important aspect of the ocean is that it stores a much larger quantity of energy than the atmo­
sphere. This is because of both its larger heat capacity and its much greater density. The verticalstructure of the ocean can be divided into two layers, which differ in the scale of their interaction
with the overlying atmosphere. The lower layer comprises the cold deep-water sphere, making
up 80% of the ocean volume. The upper layer, which has closest contact with the atmosphere, isthe seasonal boundary layer, a mixed water layer extending down only about 100 m in the tropics
but several kilometres in Polar Regions. The seasonal boundary layer alone stores approximately
30 times as much heat as the atmosphere.

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Dynamics – ocean circulation

The hydrosphere, like the atmosphere, is always in motion. The ocean temperature and salinity affectits motion. Cold water is dense and tends to sink. Sea salt is left in the water during the formationof sea ice in the Polar Regions, resulting in an increased salinity of the ocean. This cold and saline
water is particularly dense and therefore tends to sink. This takes place in certain regions, such as in
the North Atlantic, and is a driving force for the thermohaline circulation (Figure 2-5).
The transfer of heat between the surface layer and deeper layers of the ocean takes place through different mechanisms: i) diffusion or turbulent mixing which transfers heat from warmer layers to coldlayers, ii) convection which happens when surface waters become cold and denser than the water below
and thereby switch places often resulting in transferring of heat upwards, and iii) advection which isa large-scale flow transferring water from one place to another.
Advection of heat in the oceans often occurs in the form of currents that move the warm waters in thetropics toward the poles, and colder water from the Polar Regions toward the tropics. These currents
exist on the surface and at great depths of up to about 4 km. The currents are driven by the winds, bythermoclines and by the rotation of the Earth. In contrast to the atmospheric circulation and winds,
the ocean circulation and currents are more stable and reacts much more slowly to changes in external
forcing than the atmosphere. In the Polar Regions, cold water sinks and spreads throughout the oceans,causing the deep ocean to be cold. In mid latitudes, this is balanced by the upwelling of relatively warm
water. Horizontal variations in the density of the water play an important role in driving this large-scale
overturning, and in this context evaporation of water from the ocean surface is of importance, since
it increases the salinity and density of the remaining surface water.
There are linkages between anomalies in sea-surface temperatures and atmospheric long-wave circulation patterns. Shifts in sea-surface temperature anomaly patterns may therefore alter the prevailingcirculation and climate. The atmosphere-ocean system involves a two-way interaction. Atmospheric
impacts on the ocean are relatively short-term, whereas the oceans’ effect on the atmosphere isrelatively long-term because of the great thermal stability of the ocean.

Figure 2-5. The thermohaline circulation. From /Broecker 1991/.

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An important component of the system that is thought to be involved in abrupt climate change isthe ocean circulation (Figure 2-5). Here, the key aspect is the thermohaline circulation in the North
Atlantic, though other aspects of the circulation, such as wind-driven components, may also be ofimportance.
The global thermohaline circulation consists of cooling-induced deep convection; brine rejection andsinking at high latitudes; upwelling at lower latitudes; and the horizontal currents feeding the verticalflows. In the North Atlantic, where much of the deep sinking occurs, the thermohaline circulation is
responsible for the unusually strong northward heat transport. The climate in Scandinavia is in thisway to a large extent affected by the presence of the North Atlantic Drift, a northern part of the Gulf
Stream. The heat transfer by this ocean current gives a significantly warmer climate in Fennoscandia
than if this current was to be absent or reduced.

2.1.3 The cryosphere
The cryosphere comprises all the ice on Earth, that is, snow, sea ice, glaciers, ice sheets and theice in frozen ground. Areas that are snow or ice covered throughout the year include for example
Antarctica, Greenland, parts of the Arctic Ocean, and the high mountain ranges throughout theworld. Snow and ice are important for the climate system since they have a high albedo, or reflectivity. Some parts of Antarctica reflect as much as 90% of the incoming solar radiation, compared witha global average of about 30%. Without the presence of snow and ice, the global albedo would beconsiderably lower and mean global temperatures would be higher.
Exchange of water between the cryosphere and the atmosphere is considerably less than betweenhydrosphere and atmosphere. Ice has low conductivity and large thermal inertia. Both sea ice and ice
sheets also affect the oceans and ocean circulation. The large amount of freshwater stored in the ice
sheets, and variations in ice sheet volumes, play a major role when it comes to salinity distribution
and eustatic and isostatic processes that cause sea-level changes (see further Section 3.3). Throughchanges in ocean water salinity, variations in fresh water discharge from the cryosphere affect theocean thermohaline circulation.
Snow

Seasonal snow cover is the largest component of the cryosphere by area, covering up to 33% of the
Earth’s total land surface. About 98% of the total seasonal snow cover is located in the Northern
Hemisphere. Snow cover is an important climate variable because of its influence on energy andmoisture budgets. The variations in snow cover between summer and winter accounts for the large
differences in surface albedo, both on an annual basis and also between years. Snow may reflect asmuch as 90% of the incoming solar energy, whereas a snow-free surface such as soil or vegetation
may reflect only 10 to 20%. Surface temperature is highly dependent on the presence or absence ofsnow cover, and temperature trends have been linked to changes in snow cover. In addition to thealbedo effect, snow cover represents a significant heat sink during the melt period of the seasonal
cycle due to a relatively high latent melting heat. As a result, the seasonal snow cover provides
a major source of thermal inertia within the total climate system, as it takes in and releases large
amounts of energy with little or no fluctuation in temperature.
Sea ice

Sea ice is formed by the direct freezing of the water on which it floats. If the water is saline, the salt
is left in the sea water during the freezing process, making it more saline and dense, whereas sea ice
consists of freshwater.
During the period 1979–2010, sea ice in the Arctic Ocean had an average maximum extent of
15 million km2 in winter, while the average minimum extent were 7 million km2 during summer/Fetterer et al. 2002/. The corresponding numbers for the Southern Ocean around Antarctica, during
the period 1978–2007, was 18 and 3 million km2. From the numbers it is seen that the seasonal variation is considerably larger in the Antarctic.

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Sea ice is important because it regulates exchanges of heat, moisture, and salinity in the polar oceans.
It insulates the relatively warm ocean water from the cold polar atmosphere except where cracks,
or leads, in the ice allow exchange of heat and water vapour from the ocean to the atmosphere.
The number of leads determines where and how much heat and water are lost to the atmosphere, which
may affect local cloud cover and precipitation. Sea ice thickness, its spatial extent, and the fraction ofopen water within the ice pack can vary rapidly.
Sea ice acts to decouple the atmosphere and oceans, reducing the transfer of moisture and momentum,
and in this way stabilises the energy transfers within the atmosphere. The formation of sea ice inpolar regions can influence global thermohaline circulation patterns in the oceans, which greatly
influence the global climate system.
Glaciers and ice sheets

Today two large ice sheets exist on the Earth, one in Greenland and one in Antarctica. The Antarctic
ice sheet consists of the East Antarctic Ice Sheet (EAIS) and the West Antarctic Ice Sheet (WAIS). In
all continents except Australia, ice in the form of mountain glaciers, or ice caps can be found at highaltitudes. During the Last Glacial Maximum (LGM) period, around 20 kyrs BP, ice sheets covered
about one third of the total land area of the Earth (nearly 47 million km2), compared with the present
situation with ice sheets and glaciers covering about 10% (15 million km2) of the land surface.
A glacier, ice cap or ice sheet gains mass by accumulation of snow, which is gradually transformed
to ice. Mass-loss, or ablation, is mainly accomplished by melting at the surface or base, with subsequent
runoff or evaporation of the melt water. Some melt water may refreeze within the snow instead of being
lost and some snow may sublimate or be blown off the surface. Ice may also be removed by discharge
into floating ice shelves, from which it is lost by calving of icebergs and basal melting. Net accumulation occurs at higher altitude, net ablation at lower altitude; to compensate for net accumulation andablation, ice flows from higher areas to lower by internal deformation and basal sliding (see further
Section 3.1).
The Greenland and Antarctic ice sheet response to Global warming is discussed in Section 5.1. The twomost important conditions for ice surface melting involve exposure to radiation and heat exchangewith the air in contact with the ice. Radiation involves short-wave solar radiation as well as long wave
radiation from water vapour and carbon dioxide in the atmosphere. The efficiency of solar radiation inmelting is, to a large extent, influenced by the albedo of the glacier surface. If the surface is covered
with fresh snow the albedo is typically 0.6 to 0.9 and most solar radiation is reflected. If the surface
consists of glacier ice on the other hand, the albedo is 0.2 or 0.4 and melting due to the solar radiation
is much more efficient.
Heat exchange with the air at the glacier surface takes place mainly in two ways: by conduction ofheat from the air to the ice and by condensation of water vapour on the ice surface which results in
the release of latent heat. Condensation of 1g water vapour on the surface releases enough heat tomelt c. 7 g of ice. Conduction and condensation are both enhanced dramatically when windy conditions cause air turbulence close to the glacier surface.
Ice sheets have a large inertia, their response time to a change in external forcing, i.e. a climatic change,
is in the range of 100 years to 10 kyrs whereas glaciers respond much faster, typically in the range of
10 to 100 years.
Finally, changes in the basal characteristics of ice sheets can lead to changes in ice sheet dynamics. Inturn, this can release freshwater to the oceans, affecting ocean circulation patterns. Freshwater releases
from ice-related processes can be catastrophic, e.g. the breaking of an ice dam leading to the rapid
draining of the associated lake. It is believed that the draining of an ice-dammed lake in North America(Lake Agassiz) was at least partly involved in the initiation of the Younger Dryas cold climate event
during the last deglaciation.
Permafrost

Permafrost affects climate since it has an impact on vegetation and a role in the storage and releaseof carbon. Much of the Northern Hemisphere frozen ground is overlain by evergreen boreal forest.
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These boreal forests comprise both a source and a sink of carbon. The Arctic contains nearly onethird of the Earth’s stored soil carbon. If the high northern latitudes experience a significant temperature increase, the regional soils would begin to release carbon and methane into the atmosphere. This
could lead to even higher temperatures, fuelling the cycle of carbon release and temperature rise. The
presence of permafrost also has an effect on groundwater flow and composition, both in periglacial
regions and under ice sheets. For example, it may stop groundwater recharge from the surface byforming an impermeable upper layer, or it may result in freeze-out of salts in front of an advancing
permafrost freeze front (see further Sections 3.4 and 4.5.3).

2.1.4 The land surface
The appearance of the land surface is critical when it comes to transfer of energy received from the sun
to the atmosphere. Depending on the vegetation, the porosity and saturation of the soil and its thermal
properties the amount of energy that is transferred to the atmosphere can vary. The energy can be
returned as long-wave radiation, which will heat the atmosphere, or it can be used for evaporation ofwater. Soil moisture has a strong influence on the surface temperature since energy is required for its
evaporation.
Vegetation and topography determines the roughness of the land surface affecting turbulence andthe surface winds. Particles and dust from the surface are suspended into the atmosphere by thewinds and affect the radiation balance in the atmosphere. Physical and chemical processes affectcertain characteristics of the soil, such as moisture availability and water run-off, and the fluxes ofgreenhouse gases and aerosols into the atmosphere.
The presence of mountain ranges also influence global climate. North-south orientated mountain ranges
in particular have the ability to influence global atmospheric circulation patterns. In turn, climate also
affects mountain ranges, since climate governs important denudation processes, for example fluvial
or glacial erosion. These processes cause down-wearing of mountain ranges. Orogenic processes andcontinental drift, forming mountain ranges, are processes acting over time scales of millions of years.

2.1.5 The biosphere
The different ecosystems existing on Earth have significant impact on the composition of the atmospherewhen it comes to uptake and release of greenhouse gases. Photosynthesis by both terrestrial plants
and marine biota can store significant amounts of carbon from carbon dioxide, playing a central rolein the carbon cycle. Another important natural carbon sink is the oceans, where absorption of carbon
dioxide via physicochemical and biological processes takes place. Ecosystem characteristics are also ofimportance to the budgets of many other greenhouse gases, for example methane and nitrous oxide. The
exchange and storage of the trace gases are to a great extent influenced by the biosphere which in turn issensitive to climate.

2.2

Climate forcing

Climate is changing naturally on time scales of a few tens to more than 100 kyrs. These changes are
due to i) insolation changes caused by variations in the Earth orbit, ii) processes related to tectonism,
such as volcanism, and iii) alterations of the radiation emitted by the sun. In addition to the change
and variability of these natural external modifiers of the global climate system, natural internalvariability is a characteristic of the system as well. The latter has often a more regional than global
character. In addition to the natural forcing, there is also an anthropogenic forcing of the climate, one
example being the burning of fossil fuels that increases the concentration of greenhouse gases in theatmosphere.

2.2.1 Earth-orbital variations
According to the Milankovitch theory variations in the Earth’s orbit around the sun has been a main
forcing factor of the Pleistocene glacial/interglacial cycles (Figure 2-6 and 1-2). Glaciations in the past
were principally a function of variations in the Earth’s orbital parameters, and the resulting redistribution of solar radiation reaching the Earth. A key to explaining how the variation in orbital parameters
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31

Precession of equinoxes(19 and 23 kyr periodicies)

Tilt (obliquity) of Earth axis
(41 kyr periodicity)

Vega

North Star

Maximum tilt (24.5°)
Precession
Today’s tilt (23.5°)
Minimum tilt (22.5°)

EquPlane of
Earth’s orbit

Equ

ato

ato

r

r

Ax
rota is oftion

Eccentricity of Earth orbit(100 kyr periodicity)Earth

152 million km to Sun
Focus

147.5 million km to Sun

Earth

Focus (Sun)Centre of ellipse

Eliptical orbit

Figure 2-6. Earth’s orbital parameters affecting the seasonal distribution (obliquity and precession) andamount of solar radiation reaching the Earth (eccentricity). Modified from /Tarbuck and Lutgens 1999/.

can trigger ice ages is the amount of solar radiation received at Northern high latitudes during thesummer. This is critical to the growth and decay of ice sheets. When summer insolation is strong, moreradiation is absorbed at Earth’s surface at high latitudes resulting in a warmer climate. In contrast, wheninsolation is weak, less radiation is received at Earth’s surface at high latitudes, and the reduction inradiation cools the regional climate.
Latitudinal and seasonal variations in incident solar radiation due to the precession of the equinoxes(19- and 23 kyr periodicies) and variations in the tilt (obliquity) of the Earth’s axis (41-kyr periodicity) (Figure 2-6), together with important internal feed-back mechanisms in the atmosphere-oceansystem, are sufficient to drive the significant climatic changes observed on these time scales. By
comparison, the 100-kyr eccentricity periodicity is the weakest of the orbital effects. However, it isthese weaker eccentricity variations that have dominated the pattern of glacial-interglacial cycles
over the last 700 kyrs /cf. Hays et al. 1976, Imbrie et al. 1992/. Prior to this time, i.e. before the middle
Pliocene transition, the obliquity played a larger role in regulating the glacial cycles /cf. Raymo andNisancioglu 2003, Naish et al. 2009/. After the middle Pliocene transition, ice sheet volumes arelarger and deep sea water temperatures are cooler /Clark et al. 2006/. The mechanisms by which theclimate switches between different dominating orbital periodicities, and why the very weak 100-kyr
eccentricity periodicity at present is manifested as significant climate changes with growth anddecay of major ice sheets on both hemispheres, are however still not understood e.g. /Paillard 2001,
Tziperman and Gildor 2003, Rial 2004, Clark et al. 2006, Hönisch et al. 2009/. Nevertheless, it isclear that the Milankovitch theory alone is not sufficient to explain climate variability for the past
millions of years; it needs to be accompanied by major, highly non-linear, internal changes within
the climate system involving the oceans, the atmosphere and continental ice sheets.
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2.2.2 Alteration of solar radiation
The sun is the most powerful source of energy that affects the climate system. The energy output of
the sun varies over time, being correlated with sunspot activity or magnetic storms. During periodsof maximum sunspot activity, the sun emits more energy than during periods with minimum sunspot
activity. Solar variation affects the radiation balance of the Earth, but to what degree this impacts
global climate is not yet understood. Fluctuations in solar output may account for climate variations
over time scales of decades and centuries, and in the long-term on time scales up to billions of years.
2.2.3 Tectonic processes
Tectonic processes, acting over millions of years, alter the configuration of continents and oceans on
Earth and also the composition of the atmosphere by intermittent volcanic eruptions. The configurationof the continents influences the path of the ocean currents and thus has an impact on heat transport
from low to high latitudes in the oceans as well as in the global wind system. Orogenic uplift resultsin mountain ranges, which may contribute to increasing the land surface area permanently coveredby snow. Also, if there is a change in the characteristics of land and water surfaces, or in the relative
distribution of land and sea, this can affect the radiation balance and, in consequence, the climate.
Volcanism, although driven by the slow movement of the tectonic plates, occurs irregularly on much
shorter timescales. Volcanic eruptions replenish the carbon dioxide in the atmosphere, removedby the biosphere, and also emit considerable quantities of dust and aerosols. Volcanic activity cantherefore affect the energy budget and regulation of the global climate system.
Volcanic eruptions can have an immediate and profound impact on climate when fine particles suchas dust and ash together with gases are ejected into the stratosphere. Whereas volcanic modificationof the lower atmosphere is removed within days by the effects of rainfall and gravity, stratosphericmodification may remain for several years, gradually spreading to cover much of the globe. The
amounts of aerosols ejected by volcanic eruption into the atmosphere may vary on widely different
time-scales, causing natural variations in the radiative forcing. The variations can be either negativeor positive, but the climate system must react to restore the balance. Generally, a positive radiative
forcing tends to warm the surface and a negative radiative forcing tends to cool it.

2.2.4 Anthropogenic forcing
Anthropogenic forcing is a result of human activities that have an impact on the climate. Threedifferent processes are considered to have the largest effects. The first is the release of greenhousegases, mainly arising from burning of fossil fuel, which absorb and re-emit the longwave radiationfrom Earth, increasing the temperature of the atmosphere. Secondly, increased amounts of aerosolsin the atmosphere may act to either increase or decrease the temperature of the atmosphere. Thirdly,
by changing the surface characteristics of the Earth, the albedo is affected which in turn alters theradiation balance.
Climate scenarios based on global and regional climate models indicate for the 21st century a substantial rise in air temperatures, a considerable change in hydrological conditions and a rise in global
sea-level due to increased greenhouse gas levels e.g. /IPCC 2007/, see further Section 5.1 and 5.2.
The long-term effects of carbon dioxide emissions have been studied by e.g. /Archer and Ganopolski
2005, Montenegro et al. 2007, Archer and Brovkin 2008/, showing that a significant amount ofanthropogenic carbon dioxide could remain in the atmosphere for a thousand years and longer, andthat ultimate recovery occurs on time scales of tens to hundreds of thousands of years, see further
Section 5.1 and 5.2.

2.3

Climate dynamics

Many of the processes and interactions within the climate system are non-linear, meaning that there
is no simple proportional relation between cause and effect. A complex, non-linear system may
display a chaotic behaviour. The behaviour of the system is critically dependent on small changes of
initial conditions. This does not necessarily mean that the behaviour of non-linear chaotic systems is
unpredictable. The development of the daily weather is a useful example. The development of weather
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systems creating the daily weather is governed by this non-linear chaotic dynamics. This does not
exclude successful weather prediction, but the predictability of the weather becomes limited to aperiod of approximately two weeks. The climate system is also to a high degree nonlinear, but statistical characteristics of various climates may still be described and analysed by climate models.
Climate at the global, regional and local scales is forced or influenced by a wide range of factors.
These factors and their characteristic timescales are illustrated in Figure 2-7.
On a short timescale, the climate system is affected by a variety of internal feedbacks, items 6, 9, 10,
11 and 12 in Figure 2-7, counting the oceans to be an integral part of the global climate system. It isthese short-term feedbacks that contribute to the natural variability of climate on annual and decadal
timescales, in the context of which longer-term trends have to be detected and evaluated. Aperiodic
events, such as volcanic activity, can also perturb the climate on these timescales, with individual
volcanic eruptions typically affecting climate for a few years at most, though prolonged periods ofenhanced volcanism may have more profound effects.
On short timescales, the orbital characteristics of the Earth do not change substantially, so the main
factor forcing an overall trend in climate characteristics is the evolution of the atmosphere, notably
the changes in carbon dioxide and other greenhouse gas concentrations and in atmospheric aerosol
concentrations. On both short and medium timescales, internal feedbacks are known to substantiallymodify the effects of this external forcing.
On multi-millennial timescales, variations in orbital characteristics interact with changes in greenhousegas concentrations to force alterations in climate. The effects of the Milankovitch orbital variations
have been recognized for many decades. However, more recently, variations in atmospheric greenhousegas concentrations have been shown to be an important forcing factor together with orbital variations,e.g. /Cheng et al. 2009, Yin and Berger 2010/.
Imposed upon this natural variability are the changes in greenhouse gas concentrations that arise
from human activities, primarily the burning of fossil fuels. Between 1850 and 1950, the concentration of atmospheric carbon dioxide rose from a pre-industrial value of 280 ppmv to 310 ppmv,
and by 2005 the concentration was 379 ppmv /IPCC 2007/. Future increases in greenhouse gas
concentrations are estimated for a large range of emissions scenarios, and form the basis for model
simulations of future climate change /IPCC 2007/.

Characteristic time scales of climate change processes (years)
0.1

1

10

100

103

104

105

106

107

108

109

1010
1

Galactic dust/
Spiral arms

2

Evolution of the sun

3Continental drift/
Polar wandering

4

Orogeny/Isostasy

5

Orbital parameters

6

Ocean circulation

7

Evolution of theatmosphere8

Volcanic activity

9Air-sea-ice-land
feedbacks

10 Solar v ariability
Atmosphere-ocean

11 feedbacks

12 Atmosphericautovariation

Figure 2-7. Characteristic timescales of climate change processes. From /Goodess et al. 1992/.

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2.3.1 Energy budget and radiation balance
The Earth can be considered as a physical system with an energy budget that includes all gains ofincoming energy and all losses of outgoing energy. The total flux of power entering the Earth’s
atmosphere is dominated by solar radiation (99.985%). A small amount is received from geothermalenergy (0.0135%), which is produced by stored heat produced by radioactive decay leaking out of the
Earth’s interior, and tidal energy (0.002%) produced by the interaction of the Earth’s mass with thegravitational fields of other bodies such as the moon. The outgoing energy from the Earth is determinedby the albedo which on average amounts to 0.3. This means that 30% of the incident solar energy isreflected back into space, whereas 70% is absorbed by the Earth and reradiated as infrared radiation.
Of the 30% reflected energy, 6% is reflected from the atmosphere, 20% from the clouds, and 4%
is reflected from land, water and ice. The absorbed energy is eventually re-radiated mainly by theclouds and atmosphere (64%) and by the ground (6%). The radiation budget represents the balance
between incoming energy from the Sun and outgoing thermal (longwave) and reflected (shortwave)
energy from the Earth (Figure 2-3).
The various factors that influence the radiation balance can be divided into internal and externalmechanisms. Internal factors are all mechanisms affecting the atmospheric composition (volcanism,
biological activity, land use change, human activities). The main external factor is solar radiation.
External and internal factors are closely interconnected. Increased solar radiation for example resultsin higher average temperatures and higher water vapour content of the atmosphere. Water vapour
is a heat-trapping gas that absorbs infrared radiation emitted by the Earth surface. This absorption
can lead to either higher temperatures through radiation forcing or lower temperatures as a result of
increased cloud formation which in turn increases the albedo.

2.3.2 The hydrological cycle
The hydrological cycle involves the evaporation and precipitation of water, transport of water vapour
by the atmosphere, runoff from the continents to the oceans, and the net transfer of water by ocean
currents. Evaporation and precipitation account for approximately 80% of the required non-radiative
energy transfer from the Earth’s surface to the atmosphere.
The exchange of water vapour and heat between the atmosphere and oceans has a profound impact on
climate. The atmosphere and the ocean are strongly coupled to each other and the exchange between
these systems is an important way of feeding energy into the weather systems as well as the oceans.
Precipitation, condensation, and cloud formation in the atmosphere is part of the hydrological cycle,
which also affects the seas and oceans. Through precipitation and run off the salinity and thus thethermohaline circulation of the oceans can be affected.

2.3.3 The carbon cycle
Carbon dioxide is another important component that is exchanged between atmosphere and ocean
maintaining a balance. The solubility of carbon dioxide in water is temperature dependent and incold water in the polar regions carbon dioxide is dissolved and sinks into the deep ocean. Near theequator, the carbon dioxide is outgassed when the cold water upwells and warms. If there are large
quantities of sea ice present, the exchange between atmosphere and ocean is hindered.
Looking at Late Pleistocene glacial/interglacial timescales (with glacial cycles of around 100 kyrs),
there is a strong linkage between temperature changes (in atmosphere and oceans) and the amount ofcarbon dioxide in oceans and atmosphere. This interaction between temperature and carbon dioxide ison these time scales an important process that, together with e.g. orbital changes in insolation, governs
the shifts between glacials and interglacials.

2.3.4 Feedback mechanisms and climate change
Feedback mechanisms refer to processes whereby an initial change or response to an external forcing
either reinforces the initial change (positive feedback) or weakens it (negative feedback). Some ofthe processes involved in climate system feedback-mechanisms are briefly presented in Figure 2-1,
and further discussed below.

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Temperature – water vapour and cloudiness

As previously mentioned, the most important greenhouse gas is water vapour. It has a concentrationin the atmosphere that is strongly correlated to temperature and varies in concentration from 0.2% invery dry air to more than 3% in humid air. The dependence of concentration on temperature results in apositive feedback. A warming of the atmosphere allows the air to hold more water vapour resulting
in an increased absorption of heat. This enhanced greenhouse effect increases the warming, resulting ineven more water vapour in the air. A positive feedback loop is started amplifying the initial warming.
If a cooling is initiated, the feedback will result in additional cooling. However there are complications
when dealing with water vapour since the atmospheric content of water vapour affects the amount ofclouds. Although clouds can have either positive and negative feedback effects, depending on theirtype, overall an increase in cloudiness is considered to be a negative feedback, so that changes in cloud
cover characteristics compensate for more than half of the positive feedback from water vapour andalbedo changes.
Temperature – carbon dioxide

Carbon dioxide is a greenhouse gas affecting the temperature and climate of the Earth. Climate modelstudies suggest that the global mean warming from a doubling of the amount of carbon dioxide in
the atmosphere is likely to lie in the range 2°C to 4.5°C, with a most likely value of about 3°C /IPCC
2007/. An increase of greenhouse gases alters the global temperature both by increasing the amount ofheat absorbed in a cloud-free atmosphere and by trigging feedback mechanisms, both positive andnegative. One such positive feedback mechanism is the increased potential for holding more water
vapour in the atmosphere when temperature increases. Other feedbacks relate to changes in the
distribution of vegetation and polar ice sheets.
Temperature – albedo

When a cooling climate increases the extent of snow and ice, the albedo of the new snow covered
surface is increased. The high albedo causes further cooling and a positive feedback enhances the
progress of cooling. Conversely, the same feedback process amplifies climate warming.
Vegetation – albedo and precipitation

During cooling of the climate in high latitude areas, the spruce forest is replaced by tundra, increasing
the reflectivity of the land during the winter season. This causes an enhanced cooling as a positive
feedback. If the climate becomes warmer, the tundra is replaced by forest and the feedback results inmore absorbed solar radiation and a warmer climate.
If climate becomes wetter at high latitudes, the area of forests might increase. More vegetation
results in an increase of water vapour transfer back to the atmosphere. As a positive feedback,rainfall can locally increase. A change towards drier climate will work the opposite way.
Other feedback mechanisms

Increases in atmospheric turbidity (aerosol abundance) will affect the atmospheric energy budget
by increasing the scattering of incoming solar radiation. Atmospheric turbidity has been shown tobe higher during glacial episodes than in interglacials, with a consequent reduction in direct radiation
reaching the Earth’s surface. Such a situation will enhance the cooling associated with glacial periods.
Climate variability and change

As previously mentioned, climate is changing naturally on time scales of some tens to more than
100 kyrs, due to e.g. solar variability, volcanic eruptions, and insolation changes caused by variationsin the Earth’s orbit. Recent climate changes, however, are also affected by increased atmosphericconcentrations of greenhouse gases. According to /IPCC 2007/, the recent rise in greenhouse gases
is mainly due to human activities. Climate scenarios based on global and regional climate models
indicate for the 21st century a substantial rise in air temperatures, a considerable change in hydrological conditions, and a rise in global sea-level due to increased greenhouse gas concentrations in theatmo­sphere /IPCC 2007/. The impact of such changes is expected to exhibit distinct and important
geographical differences.

36

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The future climate will be shaped by both anthropogenic and natural forcing, as well as by unforced
internal variability. The scenarios for the anthropogenic climate forcing rely on assumptions about
future human activities as a result of societal choices, including future population changes andeconomic development. These scenarios for future anthropogenic climate forcing have been used toforce climate models, in order to develop scenarios for the future climate changes towards the endof the current century e.g. /IPCC 2007, Kjellström et al. 2009b/. Inferences on the internal climate
variability can be made both from past data and by means of climate modelling. Scenarios for the
future natural climate forcing are, however, difficult to construct as the relevant mechanisms are
not fully understood in a predictive sense. We can, nevertheless, estimate the possible magnitude of
future natural change, based on inference drawn from natural climate changes in the past e.g. /Moberg
et al. 2006, Kjellström et al. 2009b/.
For recent ideas on how anthropogenic forcing may modify climate today and in the future, see/IPCC 2007/ and references therein. For further descriptions of future cases of global warming,see Section 5.1 and 5.2, including specific climate modelling results relevant to the Forsmark region
presented and discussed in Section 5.1.7.

2.4

Climate in Sweden and Forsmark

2.4.1 Climate in Sweden
The following description of the present climate in Sweden is from /Ministry of the Environment
2001/. Sweden is located in the northerly west wind belt, an area where the prevailing winds comefrom the south and west. The North Atlantic Drift and the numerous areas of low pressure produce a
climate with winters that are 20–30°C warmer than at corresponding latitudes in Siberia and Canada.
The precipitation brought by the frequent low pressures gives fairly plentiful rain and snow, although
there is some rain shadow effect east of the Norwegian mountains.
According to the most frequently cited climate classification system (Köppen), Sweden has a temperate, moist climate with year-round precipitation. Along the coasts of southern Sweden, the climate
is warm-temperate, with a natural cover of deciduous forest. The climate in the rest of the country is
cool temperate, the predominate vegetation being coniferous forest. Tundra conditions prevail in themountains. The battle between areas of warm and cold air along the polar front, and Sweden’s location between the Atlantic in the west and the largest continental mass on Earth to the east, results indramatic changes in the weather, particularly in winter. Often, a change in wind direction will suffice
for icy Siberian conditions to be replaced by mild air from the Atlantic.
Summer temperatures are largely governed by altitude, and to a lesser extent by latitude. Thus themean temperature in July is 15–16°C along the entire coast. The mean temperature in summer drops
by 0.6°C with every 100 m of altitude. Even though there is little difference in temperature between
southern and northern Sweden in high summer, summer itself (defined as the time of the year when themean diurnal temperature is above 10°C) is much longer in the south than in the north. For example, insouthernmost Sweden, summer lasts for five months, compared with three in the northernmost region.
The turn of the seasons in spring and autumn, when the mean daily temperature is between 0 and 10°C,
is also much shorter in the north. So in Lapland in northernmost Sweden, winter lasts for just over half
the year, whereas the coast of Skåne, in the far south, only has winter for a few weeks. The vegetation
growing season, defined as the part of the year when the mean diurnal temperature is over 5°C, varies
considerably over the country. It lasts for between 210 and 220 days in southernmost Sweden (westernand southern Skåne and the coast of Halland), but is only half as long in the far north.
Local conditions such as topography and proximity to the sea or large lakes influence the climate
locally. The mean temperature in January can be extremely low in valleys with open terrain in inland
areas of northern Sweden (−15 to −17°C). Elsewhere in northern Sweden the January mean temperature is generally between −9 and −14°C, except along the coast in the south of the region where, as inmuch of the central inland region, the mean January temperature is −5 to −8°C. In the southern andeastern part of central Sweden, the mean temperature is −3 to −5°C in January, while it is −1 to −2°C
in southern coastal areas owing to the ameliorating effect of the nearby open sea. The temperature canvary a great deal, from approximately −50 to almost +40°C. The lowest recorded temperature is −53°C,
recorded at two locations in northern Sweden. Elsewhere in Sweden the coldest recorded temperaturesare −30 to −40°C, except along some parts of the southern coast, where it has never fallen below
TR-10-49

37

−25 to −30°C. The highest recorded temperatures display much less geographical variation than thelowest; in southern and central Sweden and along the northern coast, the records are between 34 and
36°C. Very occasionally, temperatures also rise above 30°C in other parts of the country. Over much of
Sweden annual precipitation is between 600 and 800 mm. Annual precipitation in the mountains, most
exposed to westerly winds in northern Sweden (western Lapland and Jämtland), is between 1,500 and
2,000 mm. On the western slopes of the southern uplands, maximum annual precipitation is 1,300 mm.
The Abisko area in northernmost Sweden has least precipitation, approximately 450 mm per year. This
area lies in the rain shadow of the mountains to the west.
In more or less the entire country, precipitation is heaviest during July–November. Most precipitation
falls along fronts as areas of low pressure move across the country. But several weeks may sometimes
go by in spring and early summer without any rain. Most of Sweden usually has a snow cover inwinter. In the mountains of Lapland, the ground has a snow cover for an average of 225–250 daysa year. Most of the rest of northern Sweden is covered in snow for more than 150 days a year. In
central Sweden and upland areas of the south, there is a snow cover on average between 100 and
150 days each winter. In the rest of southern Sweden, there is a snow cover for between 50 and
100 days, except along the west coast and the far south, where snow lies for less than 50 days each
winter. The maximum snow depth averages more than 60 cm throughout almost all of northern Sweden;
the mountains generally have more than a metre of snow. Air pressure distribution over the European
continent causes winds from south and west to predominate. However, winds from other directions
are fairly common because of the numerous areas of low pressure arriving from the west, and the circulation of winds around them. Occasionally, lows reaching Sweden develop into storms so intensethat winds reach hurricane force along the coast and above the tree line. For a more detailed account of
the present climate in Sweden, see /Raab and Vedin 1995/.
An overview of climate proxy data and what they can tell us about climate in Sweden during the last
millennium, or in some cases even longer back in time is presented in /Moberg et al. 2006/. That report
also provides more detailed climate information from Swedish instrumental data for the last 260 years.
A tentative evolution of climate in Sweden for the last millennium, as simulated by a climate model,
is also provided. Taken together, these three basic types of information are used to illustrate possible
Swedish climate variations during the last millennium. /Moberg et al. 2006/ provides a benchmark forthe likely range within which our regional climate has varied naturally during the last millennium.

2.4.2 Climate at Forsmark
The climate in the Forsmark region has typical values for the climate on the Swedish east coast(Table 2-1), with a mean annual air temperature of +5°C and an annual mean precipitation of
559 mm. Over the last few years (2004–2006), a time series of meteorological observations madespecifically at the Forsmark site showed that the annual mean air temperature for this short period
was +7°C and the annual mean precipitation 546 mm /Löfgren 2010/. The differences compared withthe climatological data from SMHI (Table 2-1) are probably due to a combination of the following
factors i) the short observational period at the Forsmark site, ii) local spatial variations in climate,iii) climate change over time.
The study by /Kjellström et al. 2009b/ exemplifies climates in Sweden and in the Forsmark areaby modelling possible future and past climate situations within a 100 kyr time perspective, see
Sections 4.3.2, 4.3.3 and 5.1.7. For cases of periglacial conditions (colder and drier), as well as for
future conditions dominated by global warming (warmer, wetter), the outputs from the regional climate model simulation are used to illustrate how air temperature, precipitation, runoff and evaporationin the Forsmark region may be under such climate conditions. In addition to the modelling studies,
the climate of Fennoscandia has also been studied by analyses of geological climate archives, e.g.
/Helmens 2009a/. Descriptions of these studies are given in Section 4.2.
Table 2-1. Climate data for the Forsmark region (mean values for the period 1960–1990) /SMHI
2001/. The precipitation was measured at Östhammar whereas the air temperature was measured
at Risinge.
Site

Mean annual
precipitation (mm/yr)

Mean annual air
temperature (°C)

Mean summer
Mean wintertemperature, JJA (°C) temperature, DJF (°C)

Forsmark

576

+5.0

+14.9

38

−4.3

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3

Climate and climate-related issues

3.1

Ice sheet dynamics

3.1.1 Overview/general description
Glaciers and ice sheets may form in climate regions where, in places, the winter snow precipitation
is not completely melted away during summer. A glacier is an ice mass that has been formed by successive local accumulation of snow, with ice movement due to ice deforming under its own weight.
An ice sheet is defined as a glacier that spreads out in all directions from a central dome, i.e. a large
glacier (> 50,000 km2 in area) that is not confined by the underlying topography. In reality, the flow
pattern of ice sheets is not radial from a single dome. Instead, ice often flows out from a number of
elongated ice divides, with the ice divides constituting the highest parts of the ice sheet.
Generally, ice flow velocities are moderate to slow within an ice sheet, with ice moving a few tens ofmetres per year. However, certain well-defined parts of ice sheets, so called ice streams, exhibit significantly faster ice flow. Ice streams are typically some tens of kilometres wide and several hundredsof kilometres long. Ice velocities within present-day ice streams are several hundred metres per year,
in some cases exceeding 1,000 m a–1 /cf. Joughin et al. 2004/. Since surrounding ice typically moves
considerably slower, high velocity gradients across the ice stream margins produce distinct shear
zones with heavy crevassing. Because of the high ice flux in ice streams, these features may drain large
portions of ice sheets. Ice streams are often characterised by specific basal thermal and hydrologicalconditions, differing from those of the surrounding ice sheet.
The margin of an ice sheet may be either on land or in water. If the ice margin is positioned in thesea or in a lake, it is common that the outer part of the ice sheet is floating on the water, constituting
an ice shelf. The boundary between the floating and grounded ice is the ice sheet grounding-line.
If the basal thermal and topographical conditions are favourable, sub-glacial melt water may accumulate in topographic lows beneath the ice, forming sub-glacial lakes. These are common features under
the present Antarctic ice sheet, whereas none have so far been found beneath the Greenland ice sheet.
Mass balance

The growth and decay of ice sheets are determined by the mass balance of the ice mass. The massbalance constitutes the result of the mass gain, or accumulation, and the mass loss, or ablation,
typically averaged one year. The mass gain of an ice sheet is completely dominated by the processof snow accumulation. Most of the snow falls and accumulates during winter seasons, but snow may
also accumulate during summer. For ice sheet ablation on the other hand, a number of processes
may be involved, the two most significant being surface melting of snow and ice (if the summer
climate is warm enough) and calving of ice bergs from ice shelves (when the margin is at the sea). In
addition, mass may also be lost from the ice sheet by melting of basal ice, and locally on the surface
by sublimation of ice and snow. When a certain part of an ice sheet exhibits more accumulation thanablation during one year, that part is said to belong to the accumulation area of the ice sheet. This is
in contrast to the ablation area, where there is a net loss of mass during one year. Central parts of ice
sheets typically constitute accumulation areas, whereas in case of surface melting, lower parts of ice
sheets constitute ablation areas (Figure 3-1). The line between the accumulation- and ablation areas
is called the equilibrium line.
If the total ice sheet accumulation is greater than the total ablation during one year, the mass balanceis said to be positive, whereas the opposite case produces a negative mass balance. If accumulation
and ablation are equal, the mass balance is zero.
A positive mass balance over a number of years makes an ice sheet grow, whereas a negative massbalances reduces its size. However, changes in the size of ice sheets are very slow processes acting
over hundreds and thousands of years. The response time of an ice sheet is the time it takes for asteady-state ice sheet to come to a new steady-state condition after a climate change. Since climate isconstantly changing, ice sheets are never in true steady-state, but are constantly adjusting their sizeand shape to the prevailing climate.
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39

Ice divide
Land based ice sheet

ablationarea

accumulation
area

Grounding-line
Marine based ice sheet

accumulation area

Figure 3-1. Schematic ice sheet cross section. Panel a) shows the surface net balance distribution (bN),
panel b) shows ice velocity trajectories (vxz) as well as distribution of accumulation- and ablation areas,and panel c) shows the horisontal velocity component (ux). ELA denotes equilibrium line altitude. Modified
from /Holmlund and Jansson 2003, after Denton and Hughes 1981/.

Ice temperature. The temperature of the ice is of fundamental importance for the behaviour andcharacteristics of glaciers and ice sheets. Among other things, it has a strong effect on the movement,
dynamics and hydrology of the ice. Two types of glacier ice can be defined based on temperature;1) Temperate ice (or warm ice) with the ice temperature at the pressure melting point, and 2) Polarice (or cold ice) with the ice temperature below the pressure melting point. Cold ice is harder thantem­perate ice, and also impermeable to water unless crevasses are present.
Glaciers and ice sheets are often classified according to their thermal characteristics. In the simple
case, a glacier where all ice has a temperature at the pressure melting point throughout the year iscalled a temperate glacier or temperate ice sheet, whereas a glacier in which all the ice is below
the pressure melting point throughout the year is called a polar glacier or polar ice sheet. However,an ice sheet or glacier need not consist exclusively of temperate or polar ice. In many cases, it cancontain both ice types, and in such a case it is called polythermal.
Of particular interest is the temperature of the ice at the ice sheet bed, i.e. the basal thermal condition.
An ice sheet can be cold-based or warm-based. A cold-based ice sheet has cold basal ice, and it isfrozen to its bed. There is no free water at the bed, and no sliding of basal ice over the substrate istaking place. A warm-based ice sheet is at the pressure melting point at the bed. Free water is, in thiscase, present at the ice/bed interface, and the ice may slide over the substrate. This has importantconsequences both for ice kinematics and landform development. A polar ice sheet may be either
cold-based or warm-based. In the case of a warm-based polar ice sheet, it is typically only the lower­
most part of the ice that is at the pressure melting point, whereas most of the ice sheet consists ofpolar ice. One part of a polar ice sheet may thus be cold-based at the same time as other parts arewarm-based. This is the present case for the Greenland and Antarctic ice sheets. Warm-based andcold-based ice sheets are also called wet-bed ice sheets and dry-bed ice sheets.
Glaciers and ice sheets experience melting of basal ice where the basal ice temperature is at thepressure melting point. Heat for this melting can be added from geothermal heat flux and fromfrictional heating by internal deformation of basal ice. The thermodynamic situation at the base ofan ice sheet is determined by the thermal properties of ice. Energy can be transferred by diffusion

40

TR-10-49

along a temperature gradient in ice as in all materials. However, the solidus of the ice-water vapour
phase space has a negative slope, which means that the melting or freezing temperature is depressedwith increasing pressure by 0.09 K/Pa. As a general statement, freezing of liquid water occurs whentemperature and pressure satisfy the generalised Clapeyron equation e.g. /O’Neill and Miller 1985/:

pw

ρw

−

pi

ρi

=

L
Po
							Equation 3-1
T+
273.15ρw

where pw = water pressure, ρw = water density, pi = ice pressure, ρi = ice density, L = ice density
coefficient of latent heat of fusion, T = temperature in degrees centigrade, and P0 = osmotic pressure.
Equation 3-1 couples the effect of temperature and pressure. It is a general thermodynamic relationshipnot specific for the case of ice sheets and glaciers. However, the phase change of the ice-water system
is not only controlled by temperature and pressure. Two other factors may also be of importance; 1) the
presence of solutes in water, and 2) surface tension arising from interface curvature. Just as in the casewith an increasing pressure, an increase in solutes in liquid water also depresses the melting/freezing
point. This effect is referred to as the osmotic pressure e.g. /Padilla and Villeneuve 1992/, and it isincluded in Equation 3-1. If liquid water is present at the base of an ice sheet, and it contains solutes,
this will together with the pressure modify the ice melting point. The second factor constitutes an ice/water interfacial effect. The finer grains that a sediment has, the higher the curvature of the ice-waterinterface becomes, which in turn lowers the melting point /Hohmann 1997/. For example, in clays,liquid water has been observed at temperatures down to −10°C /O’Neill and Miller 1985/.
If the effect of phase curvature is taken into consideration the Clapeyron equation may be modified
to /Raymond and Harrison 1975/:

T =−273.15  1 1 
273.15σ iw 273.15−  pw −−Po 			Equation 3-2ρ wLL  ρi ρ w L ρ irpwhere σiw = ice-water surface energy, and rp = characteristic particle radius. Equation 3-2 is thefundamental equation for the ice-water phase transition given by /Hooke 2004/. In this equation, thefirst of the three terms describes the effect of pressure on the ice-water phase transition, the second
term describes the effect of interfacial pressure, and the third term the effect of osmotic pressure.
Equation 3-2 thus gives the complete treatment of the ice-water phase transition. Commonly only
the first term is used for calculations of the pressure melting point beneath glaciers and ice sheets,
often rewritten in glaciological literature to give a simplified expression for calculating the pressure
melting point /cf. Remy and Minster 1993/:

T =−h 								Equation 3-3
1503

where T = pressure melting point temperature (°C), and h = ice thickness (m). The effect of thelowering of the pressure melting point described above is, in the case of an ice sheet, that the melting
point is lowered by c. 2 K beneath 3 km of ice. This is very important, since the basal conditions
change drastically if the bed of an ice sheet becomes melted or frozen. This affects ice sheet flow byturning on and off basal sliding, governs if glacial erosion can take place or not, and of course has a
profound impact on basal hydrology.
To demonstrate the effect of ice sheet surface conditions on the temperature distribution within polar
ice sheets, we start with a simplified case of a steady-state ice sheet with no ice flow, corresponding
to an artificial situation at an ice divide (no horizontal flow) without any precipitation (no verticalflow). Figure 3-2a shows the vertical temperature profile through such an ice mass given a surface
temperature of −20°C and a specified geothermal heat flux at the base. The resulting temperature
profile is a straight line from the surface temperature down towards the bed. In this case the verticaltemperature profile reaches the pressure melting point temperature near the bed, resulting in a warmbased polar ice sheet.
Next we consider the same case but with snow accumulation at the ice sheet surface, which means
we are introducing a vertically downward directed ice movement. Generally, this lowers the temperatures within the ice sheet as seen in (Figure 3-2b). Due to the vertical velocity and cold surface
climate, cold ice is advected downward, while the geothermal heat warms this descending ice.
TR-10-49

41

A

B
-20

Air

-20

Air

Ice sheet

Bedrock

Bedrock

C

D-20

Bedrock

Air

-20

Air

Bedrock

Figure 3-2. Vertical temperature profiles through a 2000 m thick theoretical polar ice sheet. The dashedline denotes the pressure melting point temperature. a) No ice flow (dead-ice body), b) Vertical ice flow
only – accumulation area (reflecting only a surface accumulation rate), and c) Vertical ice flow only –ablation area, d) Horizontal and vertical ice flow – accumulation area. Grey lines show the cases from theprevious plates. Modified from /Holmlund and Jansson 2003, after Sharp 1960/.

The upper part of the ice sheet develops an almost isothermal zone, whereas the ice warms quickly
near the bed. In this example the ice sheet below this accumulation area has become cold-based.
Higher precipitation rates at the surface, i.e. higher vertical velocities, result in an increased
thick­ness of the isothermal zone and also decrease the basal temperature. Furthermore, lower air
temperatures at the ice sheet surface also decrease the ice column temperatures, and vice versa.
If we instead consider an ablation area, with net mass loss at the surface, the vertical ice movement
will be directed upwards. In this case, the upward vertical velocity produces a generally warmer icecolumn, in this example resulting in temperate conditions at the bed (Figure 3-2c).
The last case to consider is a more realistic situation, when we also have horizontal ice movement.
Figure 3-2d shows a typical situation within the accumulation area of an ice sheet, but not locateddirectly on an ice divide. The horizontal velocity component is advecting cold ice into the site, icethat was formed in higher, colder parts of the ice sheet. Compared with the situation in Figure 3-2b,
the minimum temperature is now found at some depth below the surface, since the ice at the surface,formed locally at the site, is warmer (Figure 3-2d). This positive temperature gradient near thesurface has been observed in several deep drill holes, see for example the Mirny, Century 13, andByrd 9 drill hole temperatures in (Figure 3-3). Furthermore, due to the horizontal ice movement, icein the lower part of the ice sheet is warmed by internal friction, as this is where most of the internal
deformation of the ice is taking place. In this example, the ice sheet has again become warm-based.
42

TR-10-49

Temperature (°C)
–60

–50

–40

–30

–20

–10

Devon 2Ice depth (m)

500

Mirny
7.5

0
Law Dome
1.2

Ice Stream B
2.6

Century 13

Vostok 2.5

1000

Byrd 91500

2000

Figure 3-3. Borehole temperature data from the Greenland and Antarctic ice sheet. From Modified from/Holmlund and Jansson 2003, after Paterson 1994/.

Including these processes, the resulting englacial and subglacial temperatures along a flow line are
as shown in Figure 3-4. Typical polar ice sheet accumulation rates and air temperatures are assumed.
The lowest englacial temperatures are found in the highest central parts of the ice sheet, and the highest
ice temperatures are found near the front. Basal melting is taking place in the interior part of the ice
sheet and close to the margin, with a zone of basal frozen conditions in between. A narrow zone ofbasal frozen conditions at the margin may also occur due to decreasing vertical velocity.
In nature, the topography of the landscape beneath the ice sheet is also of great importance for thebasal temperature distribution. Topographic lows are more prone to experience basal melting conditions that topographic heights. This is both due to the fact that the pressure melting point is lowered
more in depressions (due to the greater ice thickness and higher basal pressures) than over surrounding
higher terrain, and due to the lower insulating capacity of thinner ice over topographic highs. The result
is, for example, that a floor of a large valley in general is more likely to have experienced longer periodsof basal melting than surrounding elevated areas e.g. /Näslund 1997/.

3000
2000

–45

1000
–35

–25

–15

–5

0
–1000

0

0

–10

20

–20

40

–30
0

500

1000

Distance from ice divide (km)1500

Basal melt rate (mm/y)

Basal temperature (°C)

4000

Altitude (m a.s.l.)

Equilibrium line

Figure 3-4. Example of modelled ice sheet temperature distribution along a flow line. Ice temperatures are
in °C. Basal melt rates are high near the terminus because of the strain heating effect from rapid ice flowin the region around the equilibrium line. Modified from /Hooke 1977/.
TR-10-49

43

An additional important process affecting ice temperature takes place if the air temperate allows
surface melting during spring periods, for example at low elevations of an ice sheet in a warming
deglaciation climate, or during times of early ice sheet formation. After the winter period with cold
temperatures, the temperature in the upper snow/firn pack of the ice sheet is well below the freezing
point. As surface melting starts, meltwater percolates down in to the snow pack and re-freezes atsome depth. During the re-freezing process, latent heat is released (334 kJ/kg, corresponding to thelatent heat of fusion for ice) which warms the surrounding snow. As the process continues the entire
snow pack can be transformed to temperate conditions during a few weeks. This is a very efficient
process, and the result is that temperate ice instead of cold ice is formed at the location. This may
have been an important process during build-up phases of Fennoscandian ice sheets, having animportant effect on the thermal characteristics of early ice sheets /Näslund 1998/.
Ice movement and thermodynamic feed-back. The stresses induced by the mass of overlaying
ice induce deformation or strain in the ice. The resulting ice movement, often referred to as internal
deformation, is present in all glaciers (being one of the main criteria for the term glacier). The shear
stress at the base of an ice sheet is calculated from:σ = ρ gh sinαEquation 3-4

where σ is the shear stress, g is the acceleration of gravity, h is the ice thickness, and α is the ice sheet
surface slope. The most common flow law of ice describing the strain rate of ice under pressure is
Glen’s flow law /Glen 1955/:nσ Equation 3-5ε =   		
B						

where ε is the strain rate, σ is the shear stress, B is a viscosity parameter that increases as the ice
gets more difficult to deform, a parameter depending on among other things ice temperature andcrystal fabric, and n is an empirically determined constant (~3) that depends on the specific creep
process that is operating. The main effect of Glen’s flow law is that moderate increases in stress (icethickness) result in a substantial increase in strain rate. For example, a doubling of the amount ofstress results in 23, that is 8, times higher deformation rate (Figure 3-5).

s
ou

fo

de

n

tio

a
rm

sc

Vi

’s
len

w

flo

law

=3

(n

)

Plastic deformation

Strain rate

In ice sheets the horizontal ice velocity increases from the ice divides, where it is zero toward themargin. The steeper surface slope at the margin induces larger driving stresses which makes theice deformation rate high. For an ice sheet with an ablation area, i.e. an ice sheet ending on land
(Figure 3-1) the maximum horizontal velocity is at the equilibrium line, whereas in the case of amarine ice sheet margin, the maximum velocity is at the ice sheet grounding-line (Figure 3-5).

G

Shear stress

Figure 3-5. Deformation rate versus shear stress for various flow laws, including Glen’s flow law for ice.
Modified from /Holmlund and Jansson 2003/.

44

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The stress that acts on the uppermost layer of a glacier does not produce a creep or plastic deformation.
Instead the uppermost part of the ice, about 30–100 m thick depending on ice temperature, is brittle,
often resulting in typical fractures, or crevasses. Most of the internal deformation takes place near
the bed where stresses are highest. The overlaying ice moves along more or less as a uniform blockon top of the deforming ice. This is indicated as the internal deformation portion of the ice velocity
in Figure 3-6.
There is also another process that may contribute to glacier movement. If the ice sheet is wet-based,
the ice may slide over the substrate, and the substrate itself may deform, especially if the water pressure
at the bed is high. The surface ice velocity is, in this case, the sum of the component from internal
deformation and the component from basal sliding and deformation (Figure 3-6). For all practical purposes, one can assume that basal sliding does not occur if the bed is frozen. Therefore, not all glaciers,or all parts of an ice sheet, have the sliding ice flow component. However, where it is present, it may
be of great importance, for example in ice streams, which are typically warm-based. Here, the observed
surface flow may be completely dominated by basal sliding /Engelhardt and Kamb 1998/. High basal
water pressures are favourable for intense basal sliding /Engelhardt et al. 1990/, as well as smooth bed
topographies and deformable, water-saturated tills e.g. /Kamb 2001, Iverson et al. 1995/.
Within glaciers and ice sheets there are important thermodynamic feedback mechanisms, in the following exemplified by a discussion on ice streams. As seen above, the viscosity of the ice is affectedby the ice temperature, with higher ice temperatures giving more easily deformable ice. At the sametime, the internal deformation itself produces frictional heat, with higher velocities producing moreheat. This gives a positive feedback mechanism; high internal deformation rates increase the temperatureof the ice, which in turn makes the ice even easier to deform.
On the other hand, looking at the basal thermal conditions of ice streams, there may be additional
thermodynamic processes acting. A fast-flowing warm-based ice stream may drain a lot of ice fromthe ice sheet. Over time this will reduce the thickness of the ice, also over the ice stream itself. The
thinner ice insulates less well from cold surface temperatures, which leads to less melting at the bed,
and eventually also to a shift from warm-based to cold-based conditions. This reduces the velocityof the ice stream considerably and the ice flux deformational heating reduces accordingly. The ice
stream now drains much less ice than before, which in time results in increased ice thicknesses/Payne 1995/. The larger ice thickness warms the bed, which again may become wet-based andbasal sliding may start again. This process suggests that ice streams may have an inherited built-in
unstable behaviour. It has been suggested as one explanation for the cyclic Dansgaard-Oeschger
events recorded in glacial marine sediment cf. /Andrews and Barber 2002/.
Total surface velocity

Basalsliding

Internal
deformation

Steel
rod
Ice

Basal
sliding

Bedrock

Figure 3-6. Internal deformation and basal sliding demonstrated by the deforming of a steel rod from time
A to A1. Modified from /Holmlund and Jansson 2003, after Sharp 1960/.

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45

In addition, the dynamics of an ice sheet, and in particular its ice streams, and the ice sheet responseto changes in climate, are to a large degree governed by the prevailing subglacial hydrology. This
field of knowledge is in rapid advance, see further Section 3.2. One process that traditionally has notbeen incorporated in large-scale ice sheet models is the coupling of longitudinal stresses within the ice
e.g. /Pattyn 2003/, that is downstream and upstream push- and pull effects, a process of importancefor proper modelling of ice streams and grounding line features. Several recent ice sheet models have
implemented this process.

3.1.2 Controlling conditions and factors
The upper boundary of the ice sheet system is the ice sheet surface, whereas the lower boundaryis the ice sheet bed, i.e. the interface between basal ice and substrate. In accordance with ice sheet
fluctuations the lateral extent of these boundaries changes over time.
Upper boundary condition – Climate. The boundary condition at the ice sheet surface is theprevailing climate, i.e. air temperature and precipitation, including its variation over time. The
air temperature at the ice sheet surface is determined mainly by the latitude, altitude and climatechanges. The latitude at a site is fixed, but the altitude varies with the local thickness variations of
the ice sheet and associated isostatic responses. The typical air temperature pattern over an ice sheet
is with the lowest temperatures in the ice sheet interior, and higher surface temperatures closer tothe margin. If temperatures are greater than 0°C during summer, surface melting takes place on thelower parts of the ice sheet surface. This melting typically amounts to several m of ice per year, withlarge variations according to the prevailing temperature regime. Surface melt rates are lower closer
to the equilibrium line. In general, during cold stages of ice sheet growth, the amount of surface melt
water production is small compared with the amounts of water produced during ice sheet deglaciation under warmer climate conditions.
Most precipitation that falls on ice sheets falls as snow. During relatively warm climates at low ice
sheet elevations, precipitation may occur as rain. Snow accumulation and surface ablation are notdistributed evenly over the ice sheet surface. The precipitation pattern reflects ice surface elevationand degree of continentality, often giving a pattern of high accumulation rates close to the ice sheet
margin with diminishing values towards the interior. On mid-latitude ice sheets, like the former
Fennoscandian ice sheet, the precipitation pattern is strongly affected by the prevailing west-windbelt and associated low-pressure tracks. This results in an orographic effect which gives mostprecipitation on the western side of large Fennoscandian ice sheets. Typical coastal accumulation
rates of the Antarctic ice sheet are 0.3–0.6 m a–1 (water equivalents), whereas the interior parts getsless than 0.1 m a–1 in precipitation /Giovinetto and Zwally 2000/, i.e. here polar desert conditions
prevail. On an ice sheet surface, regional and local variations of great magnitude in the amount ofsnow accumulation often occur, mainly due to wind re-distribution of snow in regions close to theice sheet margin (Figure 3-7).
To initiate ice sheet growth in a non-glaciated region, local climate need to change. That can either
occur due to global climate change, towards lower local summer temperatures or higher winter precipitation rates, or, over long time scales, by tectonic uplift, or a combination of climate and tectonics.
Generally, ice sheet formation in Fennoscandia involves small alpine glaciers in the Scandinavian
mountain range that grow into a mountain-centred ice sheet, and then grow to a full-scale ice sheet/Andersen and Mangerud 1989, Lundqvist 1992, Kleman et al. 1997/. Increasing evidence suggests
very dynamic ice sheet and climate behaviour during the last glacial cycle, with ice sheet growthphases interspersed by extended periods with restricted ice coverage, see also Section 4.2.
Lower boundary condition – Ice sheet bed. The lower boundary condition of the ice sheet systemis the ice sheet bed. The substrate typically either consist of bedrock, in Fennoscandia normally
crystalline rocks, or bedrock covered by till. There are two important parameters to consider here;
1) the topography of the landscape, and 2) the amount of geothermal heat flow.
The topography of the bed is of importance for the basal boundary conditions of ice sheets, as
discussed above. Above all, an ice sheet bed with strong relief produces basal thermal conditions
with stronger lateral temperature gradients than an ice sheet bed with smooth topography. The firstcase also implies larger spatial variations in basal melt water production.

46

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3000

e

rfac

0.8

Ice

t su
shee

0.6

2000

0.4
1000

Elevation (m a.s.l.)Snow accumulation (m water equivalents / year)

1

0.2

0

0
0

200

400

600

800

1000

Figure 3-7. Accumulation rates and ice surface elevations along a profile in over the East Antarctic
ice sheet. The accumulation is high in the coastal area and low in the interior. Due to wind redistributionof snow, exceptionally high values with strong variations are found in areas of nunataks and steep surface
slope. From /Richardson 2004/.

The geothermal heat flux is also of great importance for the basal condition of ice sheets e.g. /­Wadding­ton
1987/, affecting basal ice temperatures, hydrology, ice dynamics, and the erosional capacity of the ice.
Typically, for a 3 km thick ice sheet at steady-state, a 20% error in geothermal heat flux generates a 6
K error in calculations of basal ice temperatures. This has direct implications on, for example, numericalice sheet modelling. If the geothermal heat flow is not realistic in the model setup, ice sheet models
will not produce useful data on basal melting and other characteristics. Numerical ice sheet modellingstudies have also shown that basal ice temperatures are sensitive to relatively small changes in
geothermal heat flow e.g. /Greve and Hutter 1995, Siegert and Dowdeswell 1996/.
For the modelling study presented in Section 3.1.4, /Näslund et al. 2005/ calculated a distributed,
high-resolution geothermal heat flow data set for an approximate core area of the Fennoscandian
ice sheet, and embedded this within lower-resolution data published for surrounding regions. In thefollowing, a brief overview of the geothermal heat flux calculation is given.
The geothermal heat flow, or surface heat flow density (HFD), in cratonic areas consists of two
components: (1) heat produced within the mantle and core of the Earth and (2) heat produced within
the crust. The contribution from the Earth’s interior (so-called Moho- or reduced heat flow) arises
from the cooling of the Earth and formation of a solid core, and from radiogenic heat production,e.g. /Pollack et al. 1993/. The crustal component consists of radiogenic heat production where heatis produced by the natural radioactive decay of primarily 238U, 232Th, and 40K /Furlong and Chapman
1987/. The Moho heat flow has a smooth spatial variation, possibly depending on mantle convection
cell distribution /Beardsmore and Cull 2001/, whereas the spatial variation in concentration ofradioactively decaying nuclides in the lithosphere generates a heat flow with large spatial variations.
The surface HFD can be estimated by a heat flow-heat production (Q-A) relationship of the form
Q = q0+DA0								Equation 3-6/Birch et al. 1968, Lachenbruch 1968, Beardsmore and Cull 2001/ where Q is the surface heat flow
density (or geothermal heat flow), q0 is the Moho heat flow, D represents the vertical distributionof heat-producing radionuclides in the lithosphere and A0 is the radiogenic heat production fromnear-surface rocks.
The regional HFD pattern does not correlate with gravity variations /Balling 1984/, magnetic anomalies /Riddihough 1972/, or crustal thickness /Čermák et al. 1993/. However, within the Baltic shield,
as well as in other areas with similar geological settings, there is close correlation between HFD and

TR-10-4947

regional geological units, with higher heat flow from acid (commonly granitic) areas and lower heatflow values from basic areas e.g. /Landström et al. 1979, Malmqvist et al. 1983/.
The calculation of HFD values for Sweden and Finland were based on detailed data sets fromnumerous γ-emission measurements from bedrock and till. In Sweden, airborne surveys of γ-emissions
have been carried out by the Geological Survey of Sweden (SGU), sampling data at 70 m intervals
along flight lines with 17 km separation. In order to avoid shielding effects from vegetation and lakes,
only data from exposed bedrock and till surfaces were used in the calculations. The Finnish data set isbased on radiometric γ-emission measurements of 1,054 till samples providing full spatial coverage ofthe country /Kukkonen 1989/.
The calculation of HFD is performed in several steps. First, the concentrations of 238U, 232Th, and 40K
are calculated from the γ-emission measurements, using information from detailed reference measurements over calibration plates with well known isotopic concentrations. Near-surface heat production
was then calculated from the concentrations of radionuclides. In addition, the Moho heat flow contri­
bution needs to be considered. The distributed Moho heat flow by /Artemieva and Mooney 2001/ wasadded to the near-surface heat production data set. Finally, the HFD data set was re-sampled to a gridwith 5 km resolution. To provide HFD coverage for the entire model domain, data for surrounding
areas were added from the much coarser observed global HFD data set provided by /Pollack et al.
1991/. The resulting geothermal heat flow distribution is shown in Figure 3-8.
Within the LGM ice margin, the new data set shows that the geothermal heat flow varies by a factorof as much as 2.8 (geothermal heat flow values ranging between 30 and 83 mW/m2), with an averageof 49 mW/m2. This average value is 17% higher than 42 mW/m2, the typical uniform value used in
ice sheet modelling studies of Fennoscandia. Using this new distributed data set on geothermal heat
flow, instead of a traditional uniform value of 42 mW/m2, yields a factor of 1.4 larger total basal melt
water production for the last glacial cycle /Näslund et al. 2005/. Furthermore, using the new data set

Figure 3-8. Geothermal heat flow distribution over Fennoscandia. From /Näslund et al. 2005/.

48

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in high-resolution modelling, results in increased spatial thermal gradients at the bed. This enhances
and introduces new local and regional effects on basal ice temperatures and melt rates. The results
show that regional to local variations in geothermal heat flow need to be considered for proper
identification and treatment of thermal and hydraulic bed conditions under the Fennoscandian andother similar ice sheets /Näslund et al. 2005/.
Ice properties. The exponent n in Glen’s flow law (Equation 3-5) is dependent on the active creep
process. Numerous field experiments and laboratory tests of glacier ice suggest that n should be ~3.
The typical creep process taken into account is a simple power-law creep, where the creep rate is
proportional to the stress raised to some power greater than 1, for example:ε ∝ σ 3 								Equation 3-7
On the other hand, with certain temperature, stress, and grain-size combinations, diffusional creep
instead of power-law creep could take place in ice sheets /Duval et al. 1983/, which would lower thevalue of n to less than 3. Other recent studies suggest that n should be between 1 and 2 for deformation at low stresses, low temperatures, and low cumulative strains /Alley 1992, Montagnat and Duval
2000/. However, some of these conditions are likely to be important only down to depths of a few
hundred metres in the coldest parts of ice sheets.
The viscosity constant B in Glen’s flow law is dependent on a large number of parameters. Therefore,various modifications of Glen’s flow law have been developed taking into account temperature, hydrostatic pressure, and crystal orientation. The last case is the most difficult, and it is done by introducing
an anisotropy enhancement factor. The c-axes of the ice crystals are, to a large degree, uniformly
distributed in the upper part of the ice sheet, i.e. the ice is basically isotropic. When the ice over timeis affected by the weight of the accumulating overlying ice mass, it is vertically compressed and longitudinally stretched, which results in a conical distribution of c-axes orientations. Subsequently, when
the ice is affected by the simple shear close to the ice sheet bed, the c-axes are typically re-oriented toa preferred single orientation. Development of c-axis fabrics has been obtained by re-crystallisationand ice grain rotation. Since it is known from ice core studies that the crystal orientation in ice sheets is
not isotropic at depth, a non-isotropic flow law of ice improve, for example, ice sheet modelling, e.g.
/Placidi et al. 2010/.
The ice sheet dynamics described above may affect a number of geosphere variables of importancefor a deep geological repository (Table 3-1).

3.1.3

Natural analogues/observations in nature

Palaeo-ice sheets

The mid-latitudes of the Northern Hemisphere have experienced repeated continental-scale glaciations during the Late Cenozoic. As previously mentioned, these periods are referred to as glaciations
and the warm periods between the glacials are called interglacials. The present interglacial is called theHolocene, and in north-western Europe, the last glaciation is named the Weichselian.
Palaeo-ice sheets have left geological traces in previously glaciated terrain in Fennoscandia, North
America and Siberia, as well as abundant glacio-marine traces in adjacent ocean basins. Studies ofNorth Atlantic marine sediments have shown that the first traces of eroding glaciers in Greenlandand Fennoscandia date back to between 12 and 6.6 Myr BP /Jansen and Sjøholm 1991, Fronval andJansen 1996/, with larger ice sheets present in Fennoscandia from around 2.75 Myr BP. During thelast ~700 kyrs, glaciation cycles have been about 100 kyrs long cf. /EPICA community members
2004, Lisiecki and Raymo 2005/, preceded by a period of shorter (41 kyrs) glacial cycles e.g.
/Raymo et al. 1998, McIntyre et al. 2001/, see Section 2.2.2.
Geological information on till stratigraphy, interstadial deposits, glacial landforms in loose sediments
and in bedrock have for a long time been used for making reconstructions of the Weichselian glacial
history. A summary of that work is provided in /Lokrantz and Sohlenius 2006/. Section 4.2 describes
several recent studies that have been used for the reconstruction of the Weichselian glacial history
for the SR-Site safety assessment and of climate conditions during this glacial cycle.

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49

Table 3-1. A summary of how geosphere variables are influenced by ice sheet dynamics.
Geosphere variable

Climate issue variable

Summary of influence

Ground temperature

Basal ice temperature

During periods of cold-based ice coverage, low basal ice
temperatures contribute to the formation of permafrost.
During periods of warm-based ice coverage, basal ice
temperatures at the pressure melting point contribute topermafrost degradation.

Groundwater flow

Basal thermal condition
Basal melt rate
Supply of surface melt water

If the ice sheet is cold-based no free water is available andthere will be no groundwater recharge from basal melt water.
If the ice sheet is warm-based, basal melting occurs at theice/bed interface and groundwater recharge takes place. In
addition, meltwater from the ice sheet surface will, in bothcases, be transported to the bed in frontal-near areas of theice sheet (see further Section 3.2). Groundwater recharge
and flow will be determined by the presence of the ice sheet.

Groundwater pressure

Basal thermal condition
Ice sheet thickness
Basal melt rate
Supply of surface melt water

If the ice sheet is warm-based, the water pressure at the ice/
bed interface may reach as much as the ice sheet overburdenpressure, and in certain cases more. The groundwaterpressure also depends on the melt water supply and the flow
properties of the en- and sub-glacial hydrological systems(see further Section 3.2). The groundwater pressure is also
affected by the ice load compression of the bedrock poresand fractures.

Rock stresses

Basal condition
Englacial ice temperatures
Ice sheet thickness

Rock stresses will be influenced by the ice load and thehydrostatic pressure. Independently of basal conditions there
will be an increase in vertical stresses corresponding to theice thickness. The horizontal stresses will also increase. If theice sheet is warm-based the prevailing water pressures at theice/bed interface will also alter rock stresses. The alteration ofrock stresses also depends on the duration of the ice load andthe slope of the ice sheet surface. The slope of the ice sheet
surface near the front is in turn highly dependent on englacial
ice temperatures and basal thermal conditions.

Groundwater composition Glacial melt water composition

The glacial melt water is oxygen rich. The combinationof abundant melt water supply and high water pressures
may cause injection of glacial melt water to larger depths
than for non-glacial conditions. Also, the consumption ofoxygen close to the surface may be limited due to the lackof organic matter and microbiological activity.

Various types of glacial landforms may also be used to infer information on basal thermal conditionsof palaeo-ice sheets e.g. /Lagerbäck 1988a, Kleman et al. 1997, Kleman and Hättestrand 1999,
Kleman and Glasser 2007/.
Present ice sheets

At present there are two ice sheets on Earth, the Antarctic (14 Mkm2) and Greenland (1.7 Mkm2) ice
sheets. Large portions of these quasi-stable ice sheets are more than 3 km thick. In both cases, the ice
sheet load has depressed the bed so that large portions are situated below sea-level. Offshore marine
sediments have shown traces of waxing and waning Antarctic ice sheets of continental proportions
back to 40–36 Myrs /Haq et al. 1987, Hambrey et al. 1992/, and Antarctic alpine glaciers may traceback to the Oligocene /Näslund 2001/. The present configuration of the Antarctic ice sheet is thought
to have been relatively stable for the last 15–11 Myrs /cf. Shackleton and Kennett 1975, Marchant
et al. 1993/.
The spatial patterns of basal thermal characteristics of the Antarctic and Greenland ice sheets arecomplex due to ice sheet dynamics, bed topography, and geothermal heat flux variability. For both ice
sheets, parts of the bed are cold-based whereas other parts are warm-based, as seen from modellingstudies (Figure 3-9), radar soundings and drill hole data. Ice sheet modelling has also, together withgeological field observations and remote sensing studies, showed that ice streams are prominent
features of the Antarctic ice sheet, penetrating far into the ice sheet interior (Figure 3-10). This has also
been shown by remote sensing Interferometric Synthetic Aperture Radar (InSAR) studies /Joughin andTulaczyk 2002/. The outer coastal-near parts of these ice streams are typically wet-based, whereas their
upstream parts, high in the catchment areas, often are cold-based.

50

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Figure 3-9. Two examples of modelled present-day Antarctic ice sheet basal temperatures. Temperatures
are expressed relative to the pressure melting point (pmp). Figure a) is from /Johnson 2004/ and b) from
/Huybrechts 2006/.

400
350

250

-85

200
150

Velocity (ma-1)

300

100
50

0

1000 km

0

Figure 3-10. Modelled Antarctic ice surface velocities. Fast-flowing ice streams (red/yellow/green colour)
reach far into the ice sheet interior. From /Johnson 2004/ with numerical methods described in /Staiger
et al. 2005/.

During the main Pleistocene glacial periods, the Greenland and Antarctic ice sheets had a larger extent
than at present. The major limiting factor on their maximum lateral extent is the bed topography; theAntarctic and Greenland ice sheet margins are both today located close to the coast line. The continental crust does not extend out very far from the present-day coast lines, and at the continental margin
the water depth quickly becomes deeper. When the ice sheets grow larger the grounding-line migrates
outward, resulting in a larger area for the grounded part of the ice sheet. The grounding line cannot
advance past the continental margin due to the larger water depths outside. The continental margin
thus constitutes a definite constraint on the maximum spatial size of Antarctic and Greenland ice sheets

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51

during Late Cenozoic ice sheet fluctuations. The situation was similar for part of the Fennoscandian
ice sheet. From a bed topography point of view there were no constraints for ice sheet growth from the
Scandinavian mountain range towards east (the Baltic depression is not deep enough in this context),
but, towards the west, Fennoscandian ice sheets, including the last Weichselian ice sheet, could not anddid not extend farther west than the offshore continental margin /Svendsen and Mangerud 1987, Zweck
and Huybrechts 2003/.
Although the sizes of the Antarctic and Greenland ice sheets were greater during the Late Pleistocene
glacials, the maximum thicknesses need not have been larger. On the contrary, because of atmosphericmoisture starvation during the colder climates, the interior parts of the ice sheets probably were thinnerduring the coldest parts of the glacial cycles e.g. /Huybrechts 1990/. In contrast, at the ice sheet margins
the ice sheet thickness varies considerably during a glacial cycle. Areas where the ice thickness is zero
during a warm interglacial period (i.e. at the exact margin) may experience an increase in ice thicknessof several hundred metres, up to one kilometre, during a glacial maximum ice configuration /Näslund
et al. 2000/. In turn, this means that the temporal changes in basal conditions of the ice sheets, forexample changes in basal ice temperatures from the present interglacial pattern, were highly complex.
For example, due to the changed ice configuration affecting ice sheet dynamics, some parts of the
Antarctic ice sheet were warmer at the LGM ice configuration than at present, whereas at the sametime, other parts were colder (Figure 3-11) cf. /Näslund et al. 2000/.
Recently, the coupling between the dynamic behaviour of the Greenland ice sheet and ice sheethydrological processes have been inferred from a number of studies, see Section 3.2 and 5.1.

Study area
10
0

0

1900

1800

0

Ves
t

fjel

5

on
efr

5

5

0

He

0

0
5

-400

-300

10

-600

-25 deg

-500

1700

-5

im

0

tfje

lla

5

Distance (km)

la

5

-5

-10

Distance (km)

0

1600

-200

+25 deg

Figure 3-11. Modelled difference in basal temperature between the present-day ice configuration and amaximum Pleistocene ice sheet configuration for part of the Antarctic ice sheet. For this single change inice configuration, ice sheet thermodynamics induce complex changes in basal thermal pattern; some partsof the ice sheet become warmer at the bed the same time as other parts become colder. From /Näslund
et al. 2000/.

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3.1.4Model studies

Ice sheet modelling

Thermo-mechanical ice sheet modelling is a well established and useful tool for studying a large arrayof ice sheet-specific and ice sheet-related issues. Although ice sheet models have limitations, see e.g.
/SCAR 2007/, they can be used to study characteristics and behaviour of present ice sheets, palaeo-ice
sheets, and in some cases also inferred future ice sheets. As computer capacity has increased, moresophisticated models have been able to be run at higher spatial resolutions, which in turn has made itmeaningful to compare detailed model output with various types of geological information, both formodel verification and testing of glacial-geological hypotheses e.g. /Pattyn et al. 1989, Van Tatenhove
and Huybrechts 1996, Näslund et al. 2003/.
Reconstructions of palaeo-ice sheet configurations, including ice thicknesses, may be done by
two main groups of models. One group is thermo-mechanical ice sheet models that are based on
ice physics and thermo-mechanical interactions e.g. /Huybrechts 1986, Boulton and Clarke 1990,
Fastook 1994, Ritz et al. 1996, Payne and Dongelmans 1997, Marshall and Clarke 1997/. The second
group of models are isostatic rebound models that are used to invert observations of crustal reboundand shoreline locations to estimate palaeo-ice thicknesses e.g. /Lambeck et al. 1998, Peltier 2004 /.
For the present study, UMISM (University of Maine Ice Sheet Model) /Fastook and Chapman 1989,
Fastook 1990, 1994, Fastook and Prentice 1994, Johnson 1994/ was selected for the simulation oflast glacial cycle ice sheet configurations. The UMISM code was selected since a large number ofsimulations of the last glacial cycle Fennoscandian ice sheet have been done with this model over
the years e.g. /Fastook and Holmlund 1994, Holmlund and Fastook 1995, Näslund et al. 2003/. The
UMISM was part of the European Ice Sheet Modeling Initiative (EISMINT) model intercomparison
experiment and yielded output in agreement with many other physically-based ice sheet models/­Huybrechts et al. 1996, Payne et al. 2000/, which further motivates the choice of model.
The resulting reconstruction of last glacial cycle ice sheet configurations for the Forsmark site, e.g.
Figure 4-18, shows large similarities with corresponding data for a North American last glacial cycle
scenario modelled using the Peltier model /Garisto et al. 2010/. This indicates that the selection oftype of model is not critical for the general appearance of the ice sheets reconstructed for the lastglacial cycle. In detail, however, the development of ice configurations over time in e.g. these twostudies are expected to differ for several of reasons, including that the studies deal with two different
geographical areas and indeed two different ice sheets (the Weichselian ice sheet over Eurasia andthe Laurentide ice sheet over North America). For instance, the last glacial cycle Laurentide ice sheet
is known to have been considerably larger and thicker than the Weichselian ice sheet.
The ice sheet system constitutes three main sub-systems: mass-balance, ice movement, and ice
temperature. For these sub-systems the selected ice sheet model solves differential equations describing
conservation of mass, momentum, and energy respectively. In addition, the ice sheet model needs
to contain a description of climate. A mass balance parameterisation for the spatial pattern of air
temperature and precipitation is used in UMISM. The model includes a mathematical description of
precipitation based on a number of other parameters; distance from the pole, saturation vapour pressure(function of altitude and lapse-rate), and surface slope. This is an empirical relationship developed
from the Antarctic ice sheet /Fastook and Prentice 1994/. Over a certain model domain, with a
topography described by a Digital Elevation Model (DEM), this climate description gives a spatial
pattern of air temperatures at ground level and a pattern of precipitation.
The UMISM also includes a simplified isostatic description for the behaviour of deforming bedrock
due to the weight of the modelled ice sheet configuration. This is not a full self-gravitational spherical Earth model as used in the Glacial Isostatic Adjustment (GIA) modelling in Section 3.3. Instead,
it is a hydrostatically supported elastic plate model, considered adequate for the purpose of placing
the ice sheet surface at an appropriate altitude, and hence at an appropriate air temperature, for themass balance calculations. Furthermore the UMISM also includes a subglacial hydrology model/Johnson 1994/ that is able to transport basal meltwater under the ice sheet according to prevailingpressure potentials governed by ice sheet thickness and basal topography.
A palaeo-temperature record was employed to run the ice sheet model (Figure 3-12). For thereconstruction of the Weichselian ice sheet, see below, the temperature proxy record for the last

TR-10-49

53

8

2

1

4

3

5

Ice covered area (M km2)

6

Ice volume
Ice covered area
GRIP temperature curve

5
4
3
2

8

1

4

0

0
-4
-8
-12

120

110

100

90

80

70

MarineIsotope MIS 5e MIS 5d MIS 5c MIS 5b MIS 5a
Stage

Ice sheet model calibration points
1
2

MIS 5d Herning stadial
MIS 5b Rederstall stadial

60

50Time (kyr BP)
MIS 4

3
4
5

40

30

20

MIS 2

10

0

Temp. deviation from present (°C)

Ice volume (m sea level)

7

MIS 1

MIS 4 Middle Weichselian
MIS 2 Last Glacial Maximum
Younger Dryas

Figure 3-12. GRIP proxy temperature curve and reconstructed ice covered area and ice volume for thereconstruction of the Weichselian ice sheet. Times of model calibration are shown, as well as Marine
Isotope Stages (MIS) (warm stages in red and cold stages in blue).

120 kyrs from the Greenland GRIP ice core has been used /Dansgaard et al. 1993/. The temperature
file that was used contains 50 year averages of the original time series. For a detailed description ofthis temperature data set, including its use in the ice sheet modelling and its estimated uncertainties, see
Appendix 1 and /SKB 2010a/.
In the modelling process, the variations in the temperature proxy curve are used to change the climate
pattern (temperature and precipitation) over the model domain, initiating ice sheet growth inFennoscandia. Once glaciers/ice sheets are formed, the modelled thermo-dynamic ice sheet systemstarts to evolve over time. In this way, the ice sheet model simulates the behaviour of the ice sheet as
it responds to the external forcing such as changing climate or sea-level, as well as internal dynamics,
such as ice stream variations caused by internal temperature oscillations at the bed.
The following are input parameters to the ice sheet model in the simulations of the Weichselian ice
sheet:
1. The thermodynamic properties of the ice, including flow laws, sliding laws, feedback mechanisms etc.
2. Upper boundary conditions:
Climate description (annual air temperature, and precipitation rates) from the mass balanceparameterisation.

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3. Lower boundary conditions:
A. A DEM over Fennoscandia with a moderately high spatial resolution. For the reconstructionof the Weichselian ice sheet a grid size of 50 km was used. DEM data were taken from the
ETOPO2 data base1, depicting both continental topography and bathymetry;
B. Geothermal heat flux at the bed.
4. Sea-level changes during the last 120 kyrs.
5. Proxy curve of palaeo-air temperatures during the past 120 kyrs (GRIP data).
The modelling starts 120 kyrs before present, and the model was run at a five year time step forward
in time up to time zero (present). For each time step, output data are calculated for each grid cell andgrid node, such as:
1. Ice thickness.
2. Englacial and basal ice temperatures.
3. Ice velocity.
4. Direction of ice movement.
5. Isostatic depression of crust.
6. Amount of basal melting or freeze-on of water.
Model setup. A model domain was selected to cover the maximum extent of the Weichselian ice
sheet over Fennoscandia, with a spatial DEM resolution of 50 km. The thermomechanical coupling
between ice movement and ice temperature was enabled, as well as a function by which the amount ofbasal sliding is coupled to the amount of basal meltwater present at the ice sheet bed. The n parameterin Glen’s flow law (Section 3.1.2) was set to 3, whereas the values for the viscosity parameter B differs
depending on ice temperature. Exemplified calculated values of B are seen in Figure 3-13. In the final
model run, the model was run with a five year time step and data was saved from the model simulation
at a 100 year interval.
The traditional way of specifying the basal thermal boundary condition in ice sheet modelling is to
apply a single, uniform value of geothermal heat flow for the entire model domain. The magnitudeof this geothermal heat flow value has often been set according to the general geological setting forthe area to be studied. Thermodynamic modelling of Fennoscandian ice sheets has, for instance, useda single geothermal heat flow value typical for the Pre-Cambrian shield, the most common reported
value being 42 mW/m2 e.g. /Boulton and Payne 1992, Huybrechts and T’siobbel 1995, Payne andBaldwin 1999/. Others have used higher values, around 50–55 mW/m2 e.g. /Hindmarsh et al. 1989,
Forsström et al. 2003/.
For the model reconstruction of the Weichselian ice sheet, the geothermal data set of /Näslund et al.
2005/ was used as basal boundary condition. A so-called higher-order model, taking into account longitudinal push-and-pull effects within the ice sheet was used for specific studies of ice flow behaviour inrelation to geothermal heat-flow anomalies in Sweden, see /Näslund et al. 2005/.
Model calibration. In studies in which ice sheet modelling aims at simulating palaeo-ice sheets, one
needs to calibrate the behaviour of the model against known information on ice sheet extents. The
reason for this is that we do not have full knowledge on a number of the processes and boundary
conditions involved, for example on climate back in time, or ice sheet basal processes. In the present
reconstruction of the Weichselian ice sheet, the UMISM model was calibrated by making adjustments of the proxy air temperature curve (Figure 3-12). Over the years, a large number of such simulations of the Fennoscandian ice sheet have been done with the UMISM model using this climate
curve, which has given extensive experience of how to perform the model calibration. As mentioned
above, the forcing in temperature for the ice sheet model is based on the GRIP ice core /Dansgaard
et al. 1993/, see also Appendix 1. In the ice sheet calibration process, the entire temperature curve
was shifted and amplified (Figure 3-14) so as to obtain calibrated ice sheet configurations that fitted
ETOPO2, 2001. Global Digital Elevation Model (ETOPO2) representing gridded (2 minute by 2 minute)elevation and bathymetry for the world. Data were derived from the National Geophysical Data Center(NGDC) ETOPO2 Global 2’ Elevations data set from September 2001.
1

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0

45

0

0

Viscosity parameter B (bar-yr(1/3))

60

45

60

1.80
1.75
1.70
1.65
1.60
1.55
1.50
1.45
1.40
1.35
1.30
1.25
1.20
1.15
1.10
1.05
1.00
0.95
0.90
0.85
0.80
0.75
0.70
0.65
0.60
0.55
0.50

100 kmTemp. deviation from present (°C)

Figure 3-13. Calculated flow-law viscosity parameter B at 18.5 kyrs BP (at around Last Glacial Maximum),shown as vertically integrated values along the ice column. Colder and harder ice is seen in the central partsof the ice sheet and over the Scandinavian mountain range (larger B values), while warmer and softer ice isseen closer to the ice margin (smaller B values). The white line shows the modelled ice sheet margin and theblack lines shows ice thickness contours with a 500 m interval.

GRIP original
GRIP modified

4
0
-4
-8
-12
-16

MarineIsotope
Stage

120

110

100

90

80

70

60

50

40

30

20

10

0

Time (kyr BP, GRIP time scale)MIS 5e MIS 5d MIS 5c MIS 5b MIS 5a

MIS 4

MIS 2

MIS 1

Figure 3-14. Proxy palaeo-air temperature curves from the calibration process of the UMISM ice sheet
model. The final calibrated curve that was used (blue) has larger amplitude that the original data (red), i.e.
cold periods were even colder in the calibrated curve. In the calibration process, it was chosen to always
change the entire curve in the same way. The temperature curve is discussed in detail in Appendix 1.

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as well as possible with geological observations on dated marginal positions. Model calibration wasmade for ice margins as described in /Lokrantz and Sohlenius 2006/ for MIS 5d (~110 kyrs BP),MIS 5b (~85 kyrs BP), MIS 4 (~60 kyrs BP), MIS 2 (~18 kyrs BP), and for the Younger Dryas(~12 kyrs BP) (Figure 3-12). Keeping all other input data constant in the calibration process, themodified temperature curve after model calibration results in Late and Middle Weichselian ice
sheet configurations in good agreement with dated ice-marginal positions and Early Weichselian
configurations in reasonably good agreement with geological interpretations. The characteristics andapplicability of the GRIP temperature curve for the present study is further described and discussed
in Appendix 1.
The resulting calibrated reconstruction of the Weichselian ice sheet is seen in Figures 3-12 and 3-15.
Ice extends over the model domain during the cold stadials during Marine Isotope Stages (MIS)
5d, 5b, 4, and 2. During warmer interstadials, such as MIS 3, the ice cover is more restricted. As
expected, there is a clear trend of more and more pronounced ice coverage during the stadials
through the glacial cycle, with the LGM peak occurring at around 18 kyrs BP.
Different types of output data have been extracted and post-processed from the reconstruction of theWeichselian ice sheet, for example 1) basal and surface melt rates for groundwater modelling /Jaquet
and Siegel 2006/, 2) Basal and subaerial temperatures for permafrost modelling (Section 3.4 andAppendix 1), 3) ice thickness variations for modelling of global isostatic adjustment (Section 3.3)
and crustal stresses /Lund et al. 2009/, and 4) 2D ice sheet profiles for modelling of groundwaterflow under glacial conditions /Vidstrand et al. 2010/, see also Appendix 2.
For more information on the ice sheet model, see /SKB 2010g/ and for further descriptions of model
simplifications and uncertainties, see Section 3.1.7.
MIS5c (100 kyrs BP)

MIS5b (85 kyrs BP)

00

60°

12

0

60

°
60

°

60

60

°

°
60

°

60°

MIS 2 – LGM (18 kyrs BP)

500 km

60

60

°

°

°
45

0°

°

0

Younger Dryas (12 kyrs BP)

°
45

0°

°
45

0°

°
45

0°

°

0°

0°

0°

0°

MIS3 (50 kyrs BP)

60

45

60°

MIS4 (60 kyrs BP)

0°

°

45

0°

°

45

0°

°

45

0°

60°

MIS5a (80 kyrs BP)

60

MIS5d (110 kyrs BP)

00

18

00

1

60°

00 600

00
12
0
60

24

0
80

00
12
0
60

60°

12

60°

00

18

60°

00 600

12

0°

0°

0°

0°

Figure 3-15. Selected maps of ice sheet surface topography for the reconstructed Weichselian ice sheet.
The ice sheet grows to its full LGM configuration in progressive phases of increasing maximum extent, withintervening periods of more restricted ice coverage. Contour lines show ice surface elevation with a 300 m
contour interval. All maps show present day shoreline position.

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3.1.5 Time perspective
It is generally agreed that ice sheets take longer to grow than to decay. However, our knowledge of the
final deglaciation phase is much better than of inception phases, reflecting the fact that there are more
geological traces from the last deglaciation than from early stages. Since ablation by surface melting isa more efficient process than snow accumulation, this suggests that the time for ice sheet decay should
be faster than the growth time /Paterson 1994/. In reality, this is also accentuated due to efficient ice
shelf calving processes for marine parts of ice sheets, for example calving of the Fennoscandian ice
sheet in the Gulf of Bothnia. In line with the present modelling results of the Weichselian ice sheet(Figures 3-12 and 3-15) it is generally concluded that ice sheets may build up in a series of glacialbuild-up periods, peaking in a glacial maximum close to the end of the glacial. The Fennoscandian
ice sheet seems to have exhibited such dynamic behaviour during the last glacial cycle, Section 4.2.
This has also been inferred from studies of so called Heinrich events cf. /Heinrich 1988/ recorded in
North Atlantic marine sediments, suggesting that even large ice sheet systems are dynamic features,
responding relatively quickly to rapid changes in climatic forcing.
During the last 700 kyrs, glacial cycles have typically had a duration of about 100 kyrs, with interglacials lasting around 10 to 20 kyrs e.g. /McIntyre et al. 2001, Tzedakis et al. 2009/. As noted above
and discussed in Section 4.2, the last glacial cycle was also characterised by a number of cold growthphases and intervening warmer periods with less ice, although from the geological record we still do
not know how to date or correlate all observed interstadial localities /Lokrantz and Sohlenius 2006/.
A typical cold Weichselian stadial had duration of some thousands of years, see e.g. Figure 3-12 andSection 4.3.4. The deglaciation of the Weichselian ice sheet from the LGM ice margin to the last ice
remnants in the Scandinavian mountain range took about 8 kyrs, including a number of temporal halts
in the retreat, the most pronounced one occurring during the Younger Dryas (about 13–11.5 kyrs BP).

3.1.6 Handling in the safety assessment SR-Site
The evolution of ice sheets is investigated by means of numerical modelling. The base case of themain scenario in SR-Site is constructed from a reference glacial cycle based on an ice sheet modelreconstruction of the Weichselian glaciation. The thermo-dynamic ice sheet model used is capableof simulating realistic ice sheets which are typically not in balance with the climate. Derived ice
temperatures, together with density variations with depth, control ice viscosity and ice flow. The
thermodynamic calculation accounts for vertical diffusion, vertical advection and heating caused byinternal shear. The climate input, forcing ice sheet evolution, is the mean annual air temperature at
sea-level, and its variation over time. The mass balance is determined from an empirical relationshipbased on a simple parameterisation of the ice sheet’s effect on local climate /Fastook and Prentice
1994/. Distributed temperatures over the model domain are determined from height above sea-leveland distance from the pole.
The model simulation of the Weichselian ice sheet over Fennoscandia was controlled by the proxytemperature curve from the GRIP ice core, obtained from central Greenland e.g. /Dansgaard et al.
1993/, see also Appendix A, a typical method used in ice sheet modelling of the Weichselian glacial
cycle. The geothermal heat flow and its spatial variation has been shown to be of importance forobtaining realistic modelled basal ice temperatures and basal melt rates /Waddington 1987, Näslund
et al. 2005/. Basal temperatures and basal water conditions are in turn important for the overall ice flow
and ice dynamics. A detailed dataset of geothermal heat flux values, based on national measurementsof gamma emissions in Sweden and Finland has been compiled /Näslund et al. 2005/, and has forthe first time been used as input to an ice sheet model. This eliminates previous large uncertainties
in geothermal heat flow, and associated uncertainties in modelled basal conditions of the ice sheet.
The ice sheet model was calibrated so that modelled ice sheet configurations were in agreement withdated geological information on ice-marginal positions (Section 3.1.4). For this purpose, geologicalinformation on the Weichselian glaciation history of Scandinavia was compiled by /Lokrantz andSohlenius 2006/. For further information on the ice sheet modelling, see the preceding parts ofSection 3.1.

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3.1.7

Handling of uncertainties in SR-Site

Uncertainties in mechanistic understanding

Model simplification uncertainty. The most important simplifications made in the ice sheet modelling in the model reconstruction of the Weichselian ice sheet are the following.
The temperature proxy climate curve driving the model is the GRIP Greenland temperature curve.
First of all this curve describes the climatic conditions at Greenland, but in the absence of similar
palaeoclimate data from Fennoscandia, or other closer areas, these are the best data available. With
the applied modelling approach, the Greenland proxy data generate Fennoscandian ice sheets thatare in fair agreement with geological information on LGM and Younger Dryas ice margin positions,even without model calibration, which is a justification that this approach is valid. There are also
uncertainties introduced by the conversion from oxygen isotope values in the ice core to Greenland
palaeo-air temperatures /Johnsen et al. 1995/, see also Appendix 1. The temperature proxy climate
curve is probably the biggest assumption made in the modelling experiments, and therefore thisparameter has been selected for adjustment in model calibration.
In the model, the amount of basal sliding is coupled to the thickness of the basal water layer of the ice
sheet. This is a simplification, but still a reasonable way of handling the difficult topic of basal sliding
and the coupling between glacier hydrology and ice sheet dynamics. One thing that is excluded in themodel is the possible presence of sub­glacial deformable sediments, for example in the Baltic depression, which may enhance ice velocities, especially in ice streams. The exclusion of the effects of such
sediments could, in some cases, lead to some overestimation of ice thickness.
In the UMISM model, as in many other ice sheet models, the so-called shallow ice approximation isused. This means that the horizontal velocity is only dependant on the local driving stress. Horizontal
and longitudinal englacial stresses are neglect, i.e. there are no push-and-pull effects within the ice
sheet, as in nature. This is an appropriate assumption for the major part of the ice sheet, with the
exception of frontal near parts and ice divides. In contrast, the model has not been used in such a
way that this has been a shortcoming.
In the UMISM model, horizontal advection of heat was not included in the thermodynamic treatmentof the ice. The result is that the model may overestimate basal ice temperatures to some extent. Thisis likely to be of importance only in fast-flowing situations like ice streams.
The hydrostatically supported elastic plate model included in the UMISM model is a simplification,
but sufficient for adjusting bed- and ice sheet surface elevations to obtain a reasonably good surface
mass balance.
The UMISM model does not include an ice shelf model. Instead, ice shelves are simulated by adjustinga marine calving rate directly at the dynamic grounding line.
The spatial resolution of model domain is approximately 50·50 km. This resolution may, in somecases, be too coarse to correctly depict smaller features in for example geothermal heat flow pattern.
Input data and data uncertainty

The temperature proxy climate curve. See section on Model simplification and uncertainty above,
Section 3.1.4, and Appendix 1.
Digital elevation data. The accuracy of the ETOPO2 DEM is good enough for the 50·50 km and10·10 km resolution ice sheet model simulations made.
Geothermal heat flux. The high-resolution data set on spatially varying geothermal heat flux is ofhigh quality in the context of ice sheet modelling. In the new data set, high resolution geothermal data
are available only for Sweden and Finland, not for the entire area covered by the Weichselian ice sheet.
However, this does not affect modelled basal conditions in Sweden, since basically all parts of Sweden
are located down-stream from areas with high-resolution data. There are assumptions made in thecalculation of the data set that could be refined to improve the data set in the future, for example using
a varying lithospheric thickness in the calculation of surface heat flow. Also, denser data on γ-emission
measurements from bedrock are available for parts of Sweden, which may be used for future versionsof the data set. All in all, this new type of ice sheet model input has improved ice sheet modellingsignificantly from all previous reconstructions of palaeo-ice sheets by numerical modelling.
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Sea-level. The sea-level curve used as input is derived from numerical ice sheet modelling of all
Northern Hemisphere ice sheets. The present volume of the Antarctic ice sheet was held constant.
The maximum lowering of sea-level in this data set is 100 m, somewhat less than the ~120–135 m
of global sea-level lowering at the LGM deduced from coral-reef data /Yokoyama et al. 2000/.
However, this is of minor importance since the position of the grounding-line in the western Atlantic
part of the ice sheet is determined by the bed topography (i.e. continental shelf location), and theeastern part is advancing and retreating over the Baltic Sea which does not have contact with the
Atlantic during such low global sea-levels.

3.1.8 Adequacy of references
The SR-Can Climate report /SKB 2006a/, from which some of the studies are used, has undergone
QA system handling including a factual review process. Other references used for the handling of
ice sheet dynamical processes are either peer-reviewed papers from the scientific literature or from atext book /Paterson 1994/.

3.2

Ice sheet hydrology

This part of the report summarises the knowledge on water flow in and beneath glaciers and ice sheets.
Many concepts in hydrology and hydraulics are applicable to water flow in glaciers. However, theunique situation of having the liquid phase flowing in conduits of the solid phase of the same material
is not encountered elsewhere. This situation means that the heat exchange between the phases andthe resulting phase changes also have to be accounted for in the analysis. The fact that the solidus
in the pressure-temperature dependent phase diagram of water has a negative slope provides further
complications. It means that ice can melt or freeze from both temperature and pressure variations.

3.2.1 Overview/general description
The hydrology of glaciers has been reviewed by several authors /Weertman 1972, Lang 1987,
Röthlisberger and Lang 1987, Hooke 1989, Hubbard and Nienow 1997, Fountain and Walder 1998,
Schneider 2000, Boulton et al. 2001b, Jansson et al. 2003, Hock and Jansson 2005, Hock et al.
2005/. However, most reviews concern specific topics and do not provide an overview. Therefore,/Jansson et al. 2007/ provides an overview of the theoretical knowledge on water flow in and beneathglaciers and ice sheets, and how these theories are applied in models to simulate the hydrology
of ice sheets. The study provides both a broad overview and an in-depth understanding of glacierhydrology to be applied to ice sheets. In order to describe in detail the current understanding of waterflow in conjunction with deforming ice and to provide understanding for the development of ideas
and models, emphasis has in /Jansson et al. 2007/ been put on the mathematical treatments, which
are reproduced in detail. Qualitative results corroborating theory or, perhaps more often, questioning
the simplifications made in theory, are also given. The overarching problem with our knowledge ofglacier hydrology is the gap between the local theories of processes and the general flow of waterin glaciers and ice sheets. Water is often channelized in non-stationary conduits through the ice,features which due to their minute size relative to the size of glaciers and ice sheets are difficult
to incorporate in spatially larger models. New work is emerging which attempts to bridge the gap
between the ice sheet spatial-scale modelling and the details of our process understanding, see/Jansson et al. 2007/.
The general hydrological processes, summarized by e.g. /Jansson et al. 2007/ are similar under bothglaciers and ice sheets since the driving forces are the same, differing only in magnitude and spatial
scale. /Bell 2008/ has reviewed the role of water beneath the ice sheets on their stability and hence
their mass balance since the dynamics influences fluxes of ice into the oceans. Indications also emergethat temperate basal conditions favouring sliding in the presence of water may be important for theobserved waxing and waning of past ice sheets /Bintanja and van de Wal 2008/. /Bamber et al. 2007/
have summarized the developments in our view of ice sheets going from slow behemoths respondingon time perspectives of millennia or longer to dynamically active features, capable of reactions on theorder of years to decades.

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The traditional view of the glacier hydrological system is similar to a combination of a groundwater
system and a limestone karst system of shafts and tunnels, and consists of supra-, en- and sub-glacialsystem components. Figure 3-16 shows the situation on a typical glacier in summer.
The surface can be divided into two parts, a lower part where the surface consists of solid ice, referred
to as the ablation area since a net loss of mass occurs in this area, and where water will run off by
mostly channelised surface flow and an upper part, referred to as the accumulation area because a net
mass gain occurs, where the surface consists of permeable snow or firn, snow that has survived at least
one year of melting. The snow and firn pack is porous and allows water to percolate into the glacier andaccumulate at depth as firn water bodies, equivalent to groundwater bodies whereas ice effectively isimpermeable.
On an ice sheet, the snow-covered area can be further subdivided into zones depending on their
thermal conditions, see Figure 3-16. At the centre of ice sheets, an example of which is modern
Greenland, the snowpack is cold and no surface melting occurs during any part of the year. Thisis the dry snow zone. At some point at lower elevation, surface melting can occur because of
warmer conditions caused by the atmospheric temperature lapse rate. Hence there is a zone where
percolation occurs increasingly wetting the upper part of the snow pack at lower altitude since thepotential for melt increases at lower altitude. This zone is called the percolation zone. At someaunsaturated snow
and firn
Saturated firn
supraglacial channels
and ponds

englacial pockets
and conduits
bedrocksubglacial cavities
and conduits

Accumulation areab

Dry snow line

Ablationarea
Equilibrium line

Wet snow line
Snow line

Dry snowzone

Maximum surface
height in current year

Percolation zone

Wet snow zone

Surface at theend of previous
summerSuperimposed ice
zone

Surface at theend of summer

Maximum height ofsuperimposed ice

0 ° C isotherm at theend of summer

Snow
Firn with ice layers and lenses
Superimposed ice

Figure 3-16. a) The hydrological system of a glacier consist of supra-, en- and subglacial subsystems and interacts with the substrate hydrological systems. Modified from /Holmlund and Jansson 2003, after Röthlisberger
and Lang 1987/. b) The glacier surface can be divided into different facies starting with the dry snow zone
through the percolation zone to the wet snow zone. Modified from /Benson 1961, Müller 1962, Paterson 1994/.

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elevation the melting is strong enough to completely warm the snow pack. The zone of completely
temperate snow is called the wet snow zone. These zones describe the conditions met during partsof the season. During winter most of the ice sheet will be covered by dry snow and development of
the different zones initiate as air temperature rises during spring and summer. This also means thatthe different zones start to develop at lower altitude and move upglacier as the season progresses toreach their uppermost position at the end of the melt season or when the annual temperature cycle
has reached maximum temperature. This also indicates that the surface flow system will develop
throughout the entire melt season with accompanying effects on runoff and water input to an en- andsubglacial drainage system.
Water is also generated at the base of ice sheets where the basal temperature is at the pressure melting
point. The amount of energy available for melting is given by the geothermal heat flux and thedeformational energy, so called strain-heating, in the ice produced by ice flow. The geothermal heat
flux varies spatially according to the geologic structure of the Earth’s crust /Näslund et al. 2005/
but does not vary significantly in time over the time periods under consideration for a geologic
repository for spent nuclear fuel. The energy available from strain heating varies both spatially andtemporally with the waxing and waning of the ice sheet as well as with the induction and demise of
faster flow features in the ice sheet. There is, therefore, a strong feedback between ice deformation,
heat production, and water production available for enhanced sliding beneath temperate-based
parts of an ice sheet. The cold-temperate boundary is also affected by the strain heating, which may
enlarge the areas subjected to basal melting. Typical basal melt rates may vary from a few to up to
10 mm/year, although even higher melt rates can probably occur under special circumstances where
ice flow is very rapid or geothermal heat fluxes are abnormally high e.g. /Hooke 2004, p 34/. Suchcircumstances do not, however, occur on an ice sheet-wide scale.
Water generated at the surface from melt and liquid precipitation can enter the en- and subglacialsystem through crevasses and vertical shafts, moulins. If the ice is temperate, it may also be slightly
permeable. Water that enters the glacier will flow through an englacial system for some time. Commonlywater reaches the bed within a couple of ice thicknesses in distance but englacial drainage can occur
over longer distances. Most often this is associated with over-deepened basins.
The area over which surface water can penetrate the ice sheet and reach the bed is not well known.
Remote sensing of e.g. the Greenland ice sheet indicates that surface melting occurs up to very highaltitudes but also that the melting zone varies substantially from year to year. Such conditions weresurely present on the Weichselian ice sheet, especially during its retreat phase. Hence, water is inample supply. Mechanisms whereby connections from the surface to the bed may be establishedthrough ice below the pressure melting point have been debated.
However, /Alley et al. 2005b/ have shown theoretically that water-filled crevasses may penetrateeven cold ice and thus constitute a mechanism for establishing hydraulic connection between surface
and bed, even on a thick largely cold ice sheet. Although no observations have been made of suchcrevasse-based linkages, observations of surface drainage patterns on the Greenland ice sheet indicate
the presence of an intricate pattern of surface drainage and moulins connecting to the base even tensof kilometres from the terminus. The width of the zone in which water may reach the bed remains
elusive but is likely to range between tens and several tens of kilometres. It is likely that the zones on
the waning Weichselian ice sheet were at least similar in extent. Since the waning ice sheet probably
exhibited surface profiles lower in elevation than, say, the modern quasi-steady state Greenland ice
sheet, ablation zones were likely larger simply because larger areas of the Weichselian ice sheet wereat lower elevation.
Once water enters the bed it can be transported in either fast or slow flow systems. The fast systems
are able to transmit large quantities of water at low pressure. They can be thought of as tunnel systems.
The slow systems transmit smaller quantities of water while maintaining higher pressure and holding
larger volumes of water. One example of such a system is the so-called linked cavity system /Kamb
1987/. Figure 3-17 shows an example of how a slow system may be configured. The fast systems are
self-organising and tend to occur at a spacing that is determined by flow volumes, although the mechanisms are not completely understood. The slow systems are dependent on smaller-scale roughness
features. Detailed descriptions of the subglacial hydrological systems are given in /Jansson et al. 2007/.

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Direction of ice movement

CavitymetresCavityIce

Bedrock

Figure 3-17. Linked cavity system, an example of a subglacial slow flow system. Modified from /Holmlund
and Jansson 2003, after Kamb 1987/.

Because of the seasonal variations in water input from the surface (basal contributions do not vary ona seasonal basis), the en- and subglacial drainage pathways transmitting such water also experienceseasonal variations see also /Jansson and Näslund 2009/. Conduits in ice adapt to transient hydraulicconditions by creep closure if water pressure is lower than hydrostatic ice pressure in the ice surrounding the conduit. If water fluxes increase, pressure builds to force more water through the conduits. This
causes an increase in the viscous dissipation of heat and results in melting of the conduit walls andhence an increase in flow capacity of the system. The enlargement by melting may be faster than thecreep closure depending on the transient hydraulic conditions. Hence, glacier conduit systems respondquickly to increases in water fluxes by enlarging the conduits. When water fluxes decrease, pressuresin the system may drop drastically until creep closure has equilibrated to the water flux to maintain abalance between closure and conduit wall melting.
Seasonal, as well as shorter term, variations in water influx will affect the basal water pressurese.g. /Jansson and Näslund 2009/, although it is unlikely that diurnal variations in water pressure will
significantly affect the pressure variations in an underlying groundwater aquifer because of theirhigh frequency and the need for diffusion of such variations in a porous medium.
A detailed account of the supra-, en-, and subglacial hydrological system is given in /Jansson et al. 2007/.
Groundwater composition is indirectly influenced by ice sheet hydrology since it is the hydrologicalconditions at the ice/bed interface that control the recharge of glacial melt water into the bed which
in turn will affect the composition of the groundwater.

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3.2.2Controlling conditions and factors

Climate – melt water supply

The coupling between the climate system and the hydrological system of glaciers and ice sheets isthrough the energy exchange between the atmosphere and the glacier surface. Energy arriving at theice surface will be expended by first warming the surface to the melting point (if below melting tostart with) and second to melt snow and ice at the surface. The energy balance at the glacier surface
is given by
M+ΔG = R+H – LvE+LfP’						Equation 3-8
e.g. /Paterson 1994; all terms are in W m–2/ where M is the heat used to melt snow and ice; this term
represents a gain in heat and is negative if melt water refreezes in the snow pack; ΔG is the rateof gain of heat in a vertical column from the surface to the depth at which vertical heat transfer isnegligible; H is the rate of heat transfer from the air to the surface by turbulence (H is negative if the
air is cooler than the surface); Lv is the specific latent heat of vaporisation (2.8·106 J kg–1); E is therate of evaporation from the surface (E is negative if condensation occurs); Lf is the specific latent
heat of fusion for ice (3.34·105 J kg–1); and P’ is the precipitation rate of rain (negligible if the surface
is melting but significant for warming up the snow pack from refreezing under cold conditions). R is
the net radiation, defined as R = Q(1–α)+Ii–Io, where Q is the rate of incoming solar direct and scattered radiation at the surface; α is the ratio between reflected and incident solar radiation, albedo; Ii
is the rate of incoming long-wave radiation at the surface; and Io is the rate of emission of long-wave
radiation by the surface.
The surface melting can be calculated by applying Equation 3-8 to measured data from meteorologicalstations on glaciers and ice sheets. However, for estimating conditions on palaeo-ice sheets simpler
solutions must be applied since all parameters in Equation 3-8 cannot be easily estimated. This has
led to the development of statistical approaches to melt estimation. The most successful and mostwidely used is the degree-day method approach e.g. /Hock 2005/.
The degree-day method is based on the notion that melting is a function of the time the air is at a
positive temperature. Hence degree-day methods use the integrated time of positive temperature asa proxy for the heat exchange and hence melting through empirical correlation between measured
melting and integrated time of positive temperature. In its simplest form the sum is of positive
degree-days, hence the name of the method, defined by S = Σi α i Ti where Ti is the mean daily tem­
perature and α i = 1 if Ti ≥ 0°C and α i = 0 if Ti < 0°C. The melting is associated to the degree-day
sum by a degree-day factor, λ, as M = λT. /Hock 2005/ has summarised degree-day factors whichshow significant variation depending on local conditions. The degree-day approach has also beenrefined by /Hock 1998/ to yield a temporally varying sub-daily variation by introducing the potential
direct solar radiation as a modulating factor γt to the empirical relationship, Mt = γtλT, where sub­script t denoted conditions at a specific time t.
Melt water is also generated englacially, and at the base of glaciers and ice sheets. The deformationthat occurs as ice is flowing generates heat that can melt minute amounts of water. Water production
through strain heating requires the ice to be temperate and varies proportionally to the deformation
rate. There is a positive feedback between strain heating and melting in that strain heating also warms
colder ice to the pressure melting point and thereby allows the ice to melt; water increases velocity
which in turn increases the strains. The amounts produced by strain heating are relatively small.
Estimates are difficult to make since they depend on the deformation rates which vary spatially. Basedon the typical flow regime of an ice sheet, strain heating is negligible near the ice divide and is likely
to increase towards the margins of the ice sheet.
Water is also generated subglacially through energy supplied by geothermal heat. The resulting melt
rates will be directly proportional to the geothermal heat gradient in the bed rock. As with internal
heating, geothermal heat may also first warm basal ice to its melting point resulting in subsequentmelting. /Näslund et al. 2005/ have shown the importance of geothermal heat on the basal temperature and water production beneath the past Weichselian ice sheet.
The ice sheet hydrology described above may affect a number of geosphere variables of importancefor a deep geological repository (Table 3-2).

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Table 3-2. Summary of how geosphere variables are influenced by ice sheet hydrology.
Geosphere variable

Climate issue variable

Summary of influence

Groundwater flow

Sub-glacial hydrological system

Groundwater recharge and flow will be determined by thetemperature regime, melt water production and thicknessof the ice sheet (see Section 3.1). It is the configuration of
the basal hydrological system that ultimately determines
the distribution of melt water and thus groundwater
recharge and distribution of water pressures at the ice/bed
interface determining the gradients for groundwater flow.

Groundwater pressure

Sub-glacial hydrological system
Melt water supply

The groundwater pressure depends on the pressure of
the water at the ice/bed interface, the ice sheet thickness
and the ice load compressing the subsurface (seeSection 3.1). The pressure of the water at the ice bed
interface is determined by the supply of melt water andthe conductivity and distribution of conductive features inthe sub-glacial system.

3.2.3 Natural analogues/observations in nature
Our knowledge on the hydrology of glaciers and ice sheets is largely based on observations fromvalley glaciers. However, there is no reason to believe that other processes would be present inassociation with an ice sheet, although the magnitude of the processes will be scaled proportionallyto the differences between valley glacier scale (1–100 km2) and ice sheet scale (1,000–10,000 km2
and larger) /Jansson et al. 2007, Jansson and Näslund 2009/. Given today’s climatological situationin Greenland, allowing for both ice sheet surface melting and cold, high polar surface conditions in
the ice sheet interior, the best present analogue situation for a future Fennoscandian ice sheet is theGreenland ice sheet. Most of the descriptions of ice sheet observations in nature below therefore arefrom the Greenland ice sheet.
Indirect observations of the ice sheet hydrological system

There are some studies that implicitly involve a subglacial drainage system and processes therein.
Observations of large, and from a glaciological perspective, abrupt changes in ice dynamics without
corresponding changes in for example driving stress imply changes in the basal boundary condition.
Ongoing thinning of sectors of ice sheets e.g. /Wingham et al. 2009/ may occur from changes inmass balance or from changes in ice dynamics. /Joughin et al. 1996/ observed a three-fold increasein velocity over a seven week period, of the Ryder Glacier, north Greenland, after which the velocities
decreased to the original values. Drainage of supraglacial lakes into the ice sheet could have been
responsible for this event.
/Zwally et al. 2002/ used GPS measurements at the equilibrium line near Jakobshavn to show thatthe ice sheet undergoes large annual variations in surface speed. They attributed these variations tothe seasonal variations in water influx to a basal water system through moulins and crevasses, sincevelocity variations were clearly correlated to seasonal melt rates variations. /Price et al. 2008/ usedmodelling to show that variations in ice velocity, such as observed by /Zwally et al. 2002/, need not
be forced locally but can also be the result from longitudinal coupling effects from movement further
downstream.
/Moon and Joughin 2007/ used satellite interferometry to investigate the spatial variability in ice
front position of 203 Greenland outlet glaciers. Their results indicate that the behaviour is closely
linked to climate variability, thus implying that dynamic responses are forced by climate and notsustainable unforced processes. /van de Wal et al. 2008/ recorded large velocity variation along atransect of the Russell Glacier, a land-terminating outlet glacier in east Greenland. They show annual
velocity variations varying by up to a factor of four. The increase in velocity occurs within daysof measured melt rate increase. /Shepherd et al. 2009/ used a combination of satellite imagery andcontinuous dGPS measurements to show how the Russell Glacier catchment of the West Greenland
ice sheet undergoes both temporal and spatial seasonal velocity variations as well as diurnal speedup
and uplift events.

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/Joughin et al. 2004/ observed changes in the flow speed of the Jakobshavn Isbrae where velocity wasseen to decrease in the period 1985–1992 and then accelerate between 1992–2000 with additional
acceleration up to 2003. Their study clearly indicates that the ice stream response can be complex
and include both periods of deceleration and acceleration. /Holland et al. 2008/ proposed a different
cause for the speed-up of Jakobshavns Isbrae. A sudden increase in subsurface ocean temperature
occurred in 1997 along the entire west coast of Greenland. Such warm water could have accelerated
the basal melting of the ice stream at the grounding line and hence caused it to become unstable.
Whereas this process will be able to affect outlets reaching the sea, it does not help to explain thegeneral increase in velocity observed around the perimeter of the Greenland ice sheet /Rignot and
Kanagaratnam 2006, Joughin et al. 2008/. /Joughin et al. 2008/ showed that the land-terminatingparts of the Greenland ice sheet underwent seasonal velocity changes. /Howat et al. 2007/ showedthat two of Greenland’s major outlets, the Helheim and Kangerdlugssuaq, underwent a rapid increasein flow in 2004 with a subsequent decrease in 2006 to near the pre-2004 levels. They argue that this
response is due to re-equilibrium of the terminus after a calving event.
There is thus ample evidence that ice sheet dynamics change on time periods from diurnal and up to
perhaps millennia. The variations observed on diurnal to annual scales are in perfect harmony withtheory and observations on valley glaciers where the coupling to variability in the hydraulic systemis evident. Ice sheets seem to have larger time-scale variability than typically observed on valley
glaciers. This longer period variability is due to processes that occur on scales that cannot be foundon valley glaciers and thus cannot produce similar features on the smaller glaciers.
Observations of ice sheet surface melt

The surface hydrology of the Greenland ice sheet has been studied by /Thomsen et al. 1989/ in relation
to a proglacial river near Illulissat/Jakobshavn for hydroelectric power production planning. The study
showed that the surface of the ice sheet was punctured by numerous moulins, each draining a welldefined area of the ice surface (Figure 3-18). Satellite-based imagery indicates that the surface of theeastern flank of the ice sheet is covered by perennial supraglacial lakes interconnected by streams(Figure 3-19). These lakes may experience episodic catastrophic drainage e.g. /Das et al. 2008,
Shepherd et al. 2009/.
Satellite-based radar imagery also provides a good overview of the facies distribution on the Greenland
Ice Sheet. Figure 3-20 shows the conditions from the centre of the ice sheet to the margin. The ablation
area at this section of the ice sheet is approximately 100–150 km wide and the percolation and wetsnow zones combined are 300–400 km wide. Unfortunately, it is impossible to conclude how muchof the combined percolation and wet snow zone contributes to runoff. However, it is reasonable tothink that half that width may contribute. Furthermore, in /ACIA 2005/ it was shown that the different
zones vary substantially in location from year to year. The lack of long-time series of data makes it
impossible to draw any conclusions as to what might be a typical situation.
Water generated at the glacier surface, and thus available for input to the glacier hydrological system,
comes from a combination of melting of ice and snow and from rain. Direct measurements of melt
have occurred within the framework of local projects over time. /van de Wal 1996/ and /Six et al.
2001/ have summarized some of the data originating from transects. Their data shows a potential
four-year cycle in melt with no clear trend.
The Gravity Recovery and Climate Experiment (GRACE) has provided means for establishing masschange signals for regions such as the Greenland ice sheet where mass change due to climate change
occurs. /Chen et al. 2006, Luthcke et al. 2006, Ramillien et al. 2006, Velicogna and Wahr 2006/ used
GRACE data to evaluate mass losses from the Greenland ice sheet. Their results provide a strong
independent indication that mass loss from the southern part of the Greenland ice sheet has increased
substantially, thus verifying more local studies from /Rignot et al. 1997, Howat et al. 2005, Rignot
and Kanagaratnam 2006, Luckman et al. 2006/. Much of the mass losses observed are from dynamic
responses at outlet glaciers. It is unclear if this acceleration is due to increased surface melt, andassociated basal lubrication, or if there are other critical triggering factors. It seems unlikely, however,that simultaneous response over larger lengths of the Greenland ice sheet margin should be caused by
local triggering factors. /Wouters et al. 2008/ expand on the previous GRACE investigators’ results andprovide a refined and extended data record showing continued mass loss from the ice sheet including

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Figure 3-18. Surface drainage basins on the ablation area of the Greenland Ice Sheet near Illulisat/
Jakobshavn area. Filled circles mark the location of moulins in each surface drainage basin. Modified
from /Thomsen et al. 1989/.

Figure 3-19. Radarsat-1 ScanSAR image of the ablation zone of the Greenland Ice Sheet near Kangerlussuaq,
East Greenland. The background grey scale indicates surface conditions from bare ice (white) throughsuperimposed ice (medium grey) to wet snow (dark grey). The black areas are lakes that absorb the radar
waves completely. Some lakes near the right centre have complete or partial lake ice cover, in some cases
indicated by a central floe surrounded by a black rim of open water.

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Figure 3-20. Radarsat-1 ScanSAR image over east Greenland. The dark area at the centre of the image(between approximately –48° W and –44° W) comprises the percolation and wet snow zones that arebounded by the white dry snow zone to the right and the bare ice of the ablation zone to the left. The
speckled grey zone to the extreme left in the image is bare ground outside of the ice sheet. A part of Disco
Island is visible in the upper left corner of the image.

large losses from the upper parts of the ice sheet, which was not observed in earlier studies. /Hanna
et al. 2008/ verify the results from /Velicogna and Wahr 2006/ and find that the increase in melt on
the ice sheet is coupled an increase in regional Northern Hemisphere air temperatures.
Through measurements of passive microwave emissions, /Steffen et al. 2004, Figure 6.18, p 205 in
ACIA 2005/ show how the Greenland ice sheet experiences large spatially variable melt conditions.
/Bougamont et al. 2005/ have reproduced Steffen et al.’s observations by applying a surface massbalance model to the Greenland ice sheet and forcing the model with ERA-40 re-analysis data (datafrom an analysis of meteorological observations from 1957 to 2002 produced by the European Centre
for Medium-Range Weather Forecasts together with other institutions). Their work indicates surface
melt can readily be obtained from large scale re-analysis data given model tuning through surface
meteorology data from the ice sheet. /Bhattacharya et al. 2009/ use the melt-area time series covering
the period 1979–2008 to show that the melt-area abruptly increased in 1995. This change can be related
to a sign-reversal in the North Atlantic Oscillation. The change is also manifested in higher measured
temperatures around the ice sheet.
One issue in discussing ice sheet hydrology is estimating the potential melt rates under a retreat
scenario for the past ice sheets. /Ridley et al. 2005/ performed numerical modelling experiments with
the Greenland ice sheet under warming scenarios. When subjecting the ice sheet to a 4·CO2 scenario
for 3,000 years the ice sheet almost completely disappeared. One lesson from this experiment is thatwhen more ground is exposed around the perimeter of the ice sheet, feedbacks from local summer
circulation produced from warm ground around the ice sheet can augment the melting.
The Greenland ice sheet surface mass balance is obviously variable and directly coupled to the forcing
climate variability. The water produced by melting also seems to affect the dynamics, not only on
diurnal and annual time scales but also affecting longer-term trends in mass loss of peripheral parts,
probably by influencing the dynamics of the Greenland ice sheet.
Observations of shorter-term variations of subglacial water pressure have been carried out on
many glaciers, among them Storglaciären. Figure 3-21 shows a typical water pressure record from
Storglaciären, northern Sweden. The hydrological system in this part of the glacier has been inferred
to be a distributed tunnel (fast) system. The data shows how water pressures vary very rapidly inresponse to environmental factors influencing water input, air temperature governing glacier melt,
and liquid precipitation causing rapid runoff events on the glacier surface. Such fast systems arecharacterised by rapid flow and pressures that fluctuate rapidly due to the inability of the system tomaintain high pressures. For more details on this types of observations and their interpretation, see/Jansson and Näslund 2009/.

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Temperature

10

10

5

5

0

Precipitation-50

Water level above glacier bed (m)

100
120

Borehole B94-4
depth: 151 m

80

100
60

80
60

Borehole b94-2
depth: 115 m

40
20
0
185

195

205

215

225

100

75

40

50

20

25

0

0

B94-4

B94-2

Water level (% of ice overburden pressure)

(°C)

15

August(mm/hour)July
15

Time (days after Jan 1 1994)

Figure 3-21. Water pressure record from Storglaciären showing large diurnal variations in water pressure
from near zero to near overburden pressure. Longer-term changes in the pattern of the water pressure
fluctuations can also be seen. From /Jansson and Näslund 2009/.

Observations of transfer of water from the ice sheet surface to the bed

Theories of how water can create pathways from the surface of a glacier or ice sheet, through cold
ice, to the bed have been presented by e.g. /Weertman 1973, van der Veen 1998a, b, 2007, Alley et al.
2005b/. The idea is that water filled crevasses may propagate downwards through the ice. The crucial
point for both Weertman’s and Alley et al.’s work is the continuous recharge of water to the crevasse to
keep it over-pressurized. Having crevasses form across supra-glacial streams or in surface melt ponds
is thus an excellent environment for the process to be active. In the ablation area of glaciers, water isgenerally abounding and it is not hard to meet these criteria anywhere. /Fountain et al. 2005a, b/ also
observed englacial water filled crevasses at varying depths in Storglaciären, where they appear to be
an intrinsic part of the drainage system. This study changed the notion that englaclal crevasses could
not exist under typical conditions found in glaciers.
Observations of surface lake drainage events on Greenland revitalized the theories by /Weertman
1973, Alley et al. 2005b, van der Veen 2007/. /Das et al. 2008/ observed the complete drainage ofa 5.6 km2 supra-glacial lake in ~1.5 hours into an opening crevasse on the Greenland ice sheet. They
also observed significant changes in the velocity of the ice sheet as well as local uplift during theevent. From this /Das et al. 2008/ infer that water reached the bed through 980 m of ice. Although theactual drainage pathways were not investigated or mapped, water seems at least to have gained accessto the bed near the input point. This work shows compelling indications that the process described by/Weertman 1973/ and /Alley et al. 2005b/ is realistic. /Krawczynski et al. 2009/ used a model to calculate the volume of water required to propagate crevasses through the Greenland ice sheet. They foundthat small lakes of 250–800 m in diameter are sufficient to drive crevasses through 1–1.5 km of ice.
Hence the conditions for propagating crevasses through the Greenland ice sheet are commonly met.
/Catania et al. 2008/ used common offset radar images to detect drainage pathways between the surface
and the bed of the Greenland ice sheet in the vicinity of Jakobshavn. They found that detected connections are more abundant in the ablation zone than near or above the equilibrium line. Furthermore,
the drainage features coincide with inferred regions of longitudinal extension, indicating that crevasse
formation may play a decisive role for the location of moulins also on the ice sheet. In a subsequent,ground penetrating radar-based study /Catania and Neumann 2009/ identified that the drainage

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connections between surface and bed seemed to be associated with areas of significant basal melting.
Their conclusion was thus that the drainage features were persistent. This also implies that the drainage
system to some extent is persistent in space.
The transfer of water from the surface of an ice sheet such as in Greenland, through its cold interiormass of ice, is thus possible. The mechanisms outlined by /Weertman 1973, Alley et al. 2005b, vander Veen 2007/ allows for propagation of water filled surface crevasses to the bed as long as the steadysufficient supply of water is maintained. Since melt water occurs in abundance in the ablation area
and wet snow zone during the melt season, conditions are favourable over large areas of the Greenland
Ice Sheet marginal zone.
In the case of the Greenland Ice Sheet, the thermal structure is such that the ice column is cold
but may reach melting temperatures at the bed where conditions are favourable. Internal melting
from ice deformational heat is thus not expected in the Greenland ice sheet. Deformational heat
will, however, affect the temperature gradient near the base and affect the amount of basal melting
possible beneath the ice sheet.
Observations of water generation at the ice sheet bed

Some observations of basal melt regimes beneath ice sheets have been presented. /Dahl-Jensen
et al. 2003/ used NorthGRIP data and an ice flow model to calculate geothermal heat flow along aflow line obtaining variations between 50 and 200 mW/m2. This yields basal melt rates of between
7.5 mm/year at the drill site and 11 mm/year further upstream from the NorthGRIP site. /Fahnestock
et al. 2001/ used age-depth relationships from internal layering obtained from airborne radar surveys
to calculate basal melt rates in northern Greenland. In one region, melt rates reach and exceed
100 mm/year, indicating geothermal heat fluxes of up to 970 mW/m2. Such high geothermal heat
fluxes indicate the presence of a unusual geological setting not found in the marginal areas of theGreenland ice mass. /Oswald and Gogineni 2008/ used radar echo intensity to map the occurrenceof subglacial water beneath the northern part of the Greenland ice sheet. They concluded thatsignificant portions of the bed may be melting. Basal conditions at ice core drilling sites verify theirresults e.g. /Fahnestock et al. 2001, Dahl-Jensen et al. 2003/. Oswald and Gogineni’s data show thatapproximately 17% of the total flight line lengths indicated melting conditions.
Basal melt contributes significant volumes of water to the basal drainage system. In areas of surface
influx of water to the bed during summer, the basally generated volumes are over-printed by the surface
influxes. During winter, the entire basal drainage system experiences the basal melt derived fluxes.
Basal melt is thus important in that it maintains subglacial flow throughout the year and in all areas
where the ice sheet base is at the pressure melting point.
Subglacial lakes beneath ice sheets have been known to exist since the discovery of Lake Vostok, East
Antarctica /Kapitsa et al. 1996/. Such lakes may be of substantial size; the largest lake, Lake Vostok,
is 280 km long, 50–60 km wide and more than 1,000 m deep. Numerous large lakes have since been
identified to exist beneath the Antarctic ice sheet /Siegert et al. 1996, Siegert 2005, Siegert et al. 2005/
by e.g. satellite-based remote sensing methods. Because the ice sheet is locally floating in the water ofthe lakes, the basal shear stress is locally zero which produce a near horizontal ice surface over the lake.
No subglacial lakes have so far been identified beneath the Greenland ice sheet. It is not known if
subglacial lakes may have existed beneath the former Fennoscandian ice sheets.
Subglacial drainage

It is possible that unstable hydraulic situations can be caused by periodic build-up and release of
subglacial lakes. /Alley et al. 2006/ suggested a mechanism for producing periodic outburst floods andaccompanying dynamic events. The mechanism involves growth of ice shelves onto proglacial sills
where they freeze on producing a seal. Water is thus trapped in deeper parts upstream of the sill, whichwith a growing ice sheet can become over-pressurized. If the basal temperature decreases on the sill, theover-pressurized water may escape causing an outburst flood. They suggest this process may have been
involved with the so-called Heinrich events e.g. /Hemming 2004/ seen in marine sediment cores in thenorth Atlantic and for creation of subglacial lakes such as Lake Vostok. /Evatt et al. 2006/ developed anumerical model to investigate the potential for subglacial lakes. They find that lakes should fill and70

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drain periodically, basically as a function of the filling rate. The calculations allowed Evatt et al. to predictpossible locations for subglacial lakes beneath the former Laurentide ice sheet in North America. Theirlocations agree with the positions of the palaeo-subglacial lakes proposed by /Munro-Stasiuk 2003/ and/Christoffersen et al. 2008/, however, the larger predicted lakes have not yet been identified.
Subglacial drainage beneath former ice sheets have been known to exist largely in view of the vastsystems of eskers left by the ice sheet. For instance, /Clark and Walder 1994/ discussed the distributionof eskers in terms of a subglacial geology framework. They conclude that a vast system of subglacial
tunnels or canals should exist beneath an ice sheet.
In Antarctica, e.g. /Denton and Sugden 2005, Lewis et al. 2006/ discuss observable melt water generated
geomorphological features originating from larger extents of the ice sheet. /Sawagaki and Hirakawa
1997/ observed traces of meltwater in coastal areas of Antarctica. There are indications that large
quantities of water drained through the Laurentide ice sheet during its waning phase e.g. /Barber et al.
1999, Clarke et al. 2003/. Some of the observations from Antarctica have been heavily discussed, but
the evidence is mounting that major drainage events may occur. Hence, large scale, low frequency
drainage phenomena are possible in ice sheets. In the case of the Fennoscandian ice sheet al. identified
large-scale drainage phenomena are associated with subaerial lakes dammed by ice at the margin andnot subglacial phenomena.
For further detailed information on the present knowledge on ice sheet hydrology see /Jansson et al.
2007 and Jansson 2010/.

3.2.4 Model studies
In a study made for SKB /Boulton et al. 2001b/ developed a theory of sub-glacial drainage beneathice sheets based on flow in sub-glacial tunnels (R-channels), which are supposed to be the main agentsof large-scale longitudinal melt water discharge (see Figure 3-22). The widespread glacial melt water is
supplied to the tunnels as groundwater flow. The drawdown of the piezometric pressure by the tunnels
provides the mechanism for predominantly transverse groundwater flow towards the tunnels.
The tunnels are developed from water-filled cavities formed at the ice/bed surface. The groundwaterhead elevation rises as a result of increased water supply to the bed or if the ice sheet advances overa low-conductive bed. Hence, in the theory of /Boulton et al. 2001b/, tunnel spacing is a function of
the relation between bed transmissivity and melt water input.
In the summer season the high discharge by surface melting is expected to create short-lived waterfilled voids between the large tunnels (Figure 3-23). Many of these voids will form locally beneaththe bases of water conducting moulins and crevasses. In the winter seasons such drainage routes
will be closed by internal ice flow. /Boulton et al. 2001b/ argued that the relatively long winters andshort summers representative of the climate around large ice sheets will inhibit the development ofwell-integrated summer drainage patterns, except very close to the ice sheet margin.
Winter condition

Summer condition
Groundwater flow only

Melting zone
Channelled flow

Groundwater flow only

Figure 3-22. Postulated patterns of groundwater flow in relation to the distribution of subglacial tunnels.
From /Boulton et al. 2001b/.
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Summer condition

Winter condition

pc

pw

pi

∆pw

Equilibrium water pressure

Enlarged longterm tunnels

Small ephemeralchannels and voids

Groundwater piezometric surface

Figure 3-23. Left: Transverse cross-section of a glacier showing the postulated relationships between
groundwater surface and sub-glacial tunnels during the winter period. Right: The situation during thesummer period of high discharge. Interfluve zones of high pressure, where effective pressures tend to zero,
will tend to create short-lived water filled voids. From /Boulton et al. 2001b/.

Ice sheet hydrology can be modelled in one of two main ways. Commonly the flow is represented
by simple routing routines that follow the basic ideas laid down by /Shreve 1972/ where the discrete
flow routing of water is not described, but where the water pressure potential is used to predict thedirection of flow. This type of modelling probably correctly handles the flow paths of water but not
the quantitities of water involved. Only three models exist in which the entire hydrological system orparts of it are modelled using a physically based model. In these models, the paths and quantities ofwater flowing in the system are handled. The choice of model depends on the purpose of the analysis
in question, what geometric scale is to be studied, and available data. Two models describe thecomplete hydrology of a glacier /Arnold and Sharp 2002, Flowers and Clarke 2002a, b/.
The present state of physically-based models is still crude, in that these models have been tailored to
specific glaciers or glacier types and are not necessarily easily and generally transferable. Melt waterproduction and the supra-glacial system are handled comparatively well; it is probably also the simplest
part to model since the entire system is easily observable. The en- and sub-glacial systems are more
complicated, characterised by large variability in possible drainage system configurations. Much uncertainty yet remains as to the exact geometry of drainage systems and how these systems can co-exist or
switch from one state to another, thus limiting the applicability of physically based modelling.
Modelling the physics of glacier hydrology is complex, since it involves the liquid phase, namelywater, moving through the solid phase, namely ice, at the melting temperature. Furthermore, ice isdeformable under relatively low stresses, which allows channels and conduits in glaciers to change
size and shape much more rapidly than channels eroded in rock or sediment. The glacier system isthus transient on all time scales and over all spatial scales.
The first physically based models in glaciology were made to reproduce simple observable phenomena
in glacier hydrology. /Röthlisberger 1972/ and /Shreve 1972/ presented simple numerical models(Röthlisberger: conduit flow; Shreve: water pressure potential flow) for water flow in glaciers. The
assumptions made were that the rate of melt enlargement is balanced by creep closure (Röthlisberger)
and that water pressure is in balance with ice pressure /Shreve 1972/. /Nye 1976/ and /Spring and
Hutter 1981/ presented a physically based theory of unsteady flow through intraglacial channels.
/Hooke 1984/ investigated the effects of non-steady conditions, in particular open channel flow, on the
Röthlisberger system. This was elaborated by /Kohler 1995/ who investigated the ratio of open-to-filled
channel flow in a subglacial system. Most models have assumed circular cross-section pipe flow andsemi-circular tunnels. However, discrepancies between measurements and model calculations led/Hooke et al. 1990/ to propose low broad tunnels in place of semicircular ones in order to be able
to apply Glen’s flow law of ice deformation with reasonable viscosity parameter values. Conduit
flow was found not to reproduce water pressure and fluxes observed during the surge of VariegatedGlacier, Alaska. This prompted /Kamb 1987/ to propose the linked cavity system which combined large
cross-section with small flow velocities. /Walder and Fowler 1994/ extended Röthlisberger’s model toinclude ‘canals’ cut into deformable till.

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A first approach to producing a complete physically based model for a glacier was reported in /Arnold
et al. 1998/, henceforth referred to as the Arnold model. This model has three components. A surfaceenergy sub-model calculates distributed glacier surface melt water production. A routing sub-modeltakes rain- and melt-water and either delivers the total water flux to a priori prescribed moulins,
where it enters the glacier interior or supraglacially to the ice margin. The third and final component
is a subglacial sub-model that involves a conduit system. The englacial drainage is treated as part ofthe subglacial conduit system and simulated analogous to a sewage pipe system. Subglacial conduits
are fed by a network of “drains”, which represent moulins where water can enter or overflow fromthe system. Subglacial conduits can enlarge and contract in response to changes in the rates of wall
melting and creep closure associated with changes in water inputs. In addition, the configuration of
the system can change between a distributed system and a channelised system represented by different
predefined numbers and geometries of conduits. “Distributed” links were changed to “channelised”
flow paths as the modelled snowline passed in the up-glacier direction over each moulin. From thisperspective the subglacial model is rigid, since switching between systems does not permit smoothtransitions with the bed being partially drained by each system. The model was applied to Haut Glacierd’Arolla, Switzerland, and performed well in comparison with proglacial stream discharge, but limitations are indicated by discrepancies between model outputs and field observations of subglacial waterpressure and water velocities, although the substantial features of these records could be reproduced.
/Flowers and Clarke 2002a, b/, henceforth referred to as the Flowers model, developed a different
approach to a complete physically based model. The model comprises coupled surface runoff,
englacial, subglacial, and groundwater systems. Each of the four components is represented as a twodimensional, vertically integrated layer that communicates with its neighbours through water exchange.
Melt is computed by a distributed temperature index model. Englacial hydrology is represented bydescribing a variety of bulk storage elements and allowing water transport between them in a system ofcracks. Hence, the englacial system is treated as fracture-connected crevasses and pipes. The Flowers
model was built to represent conditions at Trapridge Glacier, Canada, which is underlain by thickporous sediments. Their subglacial model, therefore, comprises flow through macroporous glacier sediments at the ice-bed interface and subsurface aquifers of buried sediment layers not directly exposedto the base of the glacier. Hence there are no explicit tunnels as in the Arnold model. The model
reproduced well diurnal cycles of subglacial water pressure as measured in boreholes.
A crevasse-like englacial network, as adopted in the Flowers model, has been observed on Storglaciären,Sweden /Fountain et al. 2005a, b/, countering previous notions of englacial water flow through few
melt-enlarged conduits. If permeable subglacial sediments are present, groundwater flow may be asignificant part of the system, as in the Flowers model, or flow may occur in channels eroded into
the subglacial sediments. On harder beds or beds of less permeable sediments such as tills, waterflow may occur in conduit systems melted into the ice or in linked cavity systems. Hence, a general
physically-based model would have to accommodate all these possibilities and even couple different
types of systems beneath different parts of a glacier. Such a model would be inherently complex.
Another issue that is not satisfactorily addressed in existing models is the time-transgressive develop­
ment of subglacial systems. In the Flowers model, this is not necessary since flow through porous
sediments does not involve significant time-dependent changes, except possible changes such as development of piping or siltation of the porous media from fines produced through sediment deformation.
In the case of the Arnold model, it is widely reported e.g. /Nienow et al. 1998/ that the subglacial drainage system changes both in sizes of conduits and in complexity of the network of channels through thecourse of a season. The subglacial system is very dynamic and it seems as if a complete and accurate
model description of this system may be distant. However, most changes in such a system occur in
response to rapid changes, both increases and decreases in water inputs, so a first-order approximation
may be to switch between a series of systems prompted by key events in the forcing.
Hence the state of glacier hydrological modelling adequately reproduces the large-scale features,such as the regional direction of flow. Finer details, such as the position of individual conduits andthe behaviour of the conduits and the water pressure within them are not completely resolved.

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3.2.5 Time perspective
The time perspective of glacier hydrology can be divided in two main categories, times scales underwhich the ice sheet undergoes significant changes in volume and extent and where the associatedhydrological conditions change with the ice sheet; and time scales under which the ice sheet can beconsidered constant in volume and extent and where the variability of the acting processes dominates.
Seen under the perspective of an entire glaciation, the hydrological system varies spatially with thesize of the ice sheet. This means that different parts of the bed are either temperate or cold and can beassociated with hydrological systems of different overall structure and parameterisation. The distribution of cold and temperate basal conditions changes in time and also in location with the waxing andwaning ice sheet. From this perspective, a possible repository site will experience changing conditionsof active hydrology under thawed bed conditions or cold frozen bed conditions. The location of large
tunnel systems also varies on this time scale. This is because, with changing ice sheet configuration,
the basal hydraulic conditions vary so that shifts in location of drainage pathways occur. How muchtime a specific location is subjected to either kind of condition is determined by the large-scale
fluctuations of the ice volume and extent.
The processes in the hydrological system of an ice sheet exhibit variations ranging from diurnal(in the case of temperature-dependent melt variations) or possibly shorter duration (in the case of
rainfall-induced variations) to seasonal variations that depend on temperature variations on an annual
basis. Since all variations in temperature-induced melting and liquid precipitation yield variationsin input to the hydrological system, they also yield variations in subglacial water pressure withoscillations on similar time scales. This means that the base of an ice sheet will exhibit variations on
a diurnal and subdaily scale superimposed on longer-term variations coupled to weekly variations
due to changing air masses of different temperature and lastly variations due to seasonal temperaturefluctuations.

3.2.6 Handling in the safety assessment SR-Site
Information from natural analogues and model studies, Section 3.2.3 and 3.2.4 has been used forgiving a description of the evolution of hydrological conditions and groundwater for glacial periodsin the reference glacial cycle (Section 4.5.5), and also for setting up boundary conditions anddiscussing results of groundwater flow modelling for glacial conditions.
3.2.7

Handling of uncertainties in SR-Site

Uncertainties in mechanistic understanding

Although the theories and empirical studies of glacier hydrology are of appropriate quality toreproduce observed phenomena at individual glaciers or sites on glaciers, there are no generally
applicable theories for the spatial and temporal variations in hydrology and hydraulic conditions.
Perhaps more importantly, the coupling between discrete drainage pathways and the continuum
mechanics of ice sheet models is not, thereby preventing integrated modelling. The main obstacle
to enhancing our understanding of the dynamics of glacier hydrology is the transient nature of thesystem and difficulties in observing its structure and behaviour. The division between subglaciallyderived and surface-derived water emanating from, say, the Greenland ice sheet is also unknown.
Almost exclusively, process studies have been made on valley glaciers. However, there is no reason
to believe that other processes would be present in association with an ice sheet, although the magnitude of the processes will be scaled proportionally to the differences between valley glacier scale andice sheet scale /Jansson et al. 2007, Jansson and Näslund 2009/.

3.2.8 Adequacy of references
The SKB reports produced for the handling of glacial hydrology /Jansson et al. 2007, Jansson and
Näslund 2009/ have undergone the SR-Site QA system handling, including a documented factual
review procedure. Also the SR-Can Climate report /SKB 2006a/, from which some of the studies areused, has undergone QA system handling including a factual review process. Other references used
for the handling of the permafrost and freezing processes are either peer-reviewed papers from thescientific literature or from a text book /Paterson 1994/.
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3.3

Isostatic adjustment and shore-level changes

In addition to the descriptions in the present section, a detailed in-depth description of the physics ofglacial isostatic adjustment (GIA), how it affects sea-level, and the methods which are employed to
study and understand these processes are presented in /Whitehouse 2009/.
The major part of Section 3.3.1 is adapted from /SKB 2006a, Section 3.3/ written by P. Whitehouse,
whereas some parts are adapted from /Lund and Näslund 2009/. Section 3.3.2 is adapted from /SKB
2006a/. Section 3.3.3 is from /Lund and Näslund 2009/, and Section 3.3.4 is adapted from /SKB 2006a/
and /Whitehouse 2009/.

3.3.1 Overview/general description
The redistribution of mass associated with the growth and decay of continental ice sheets gives rise to
major glacial loading and unloading effects over time scales of several tens of thousands of years. The
response of Earth’s crust, mantle and gravitational field is referred to as Glacial Isostatic Adjustment(GIA). For instance, during the decay of a major ice sheet, the unloading of mass results in a postglacial rebound of the crust, which continues well after the disappearance of the ice. This process iswell known from previously glaciated regions such as Canada and the United States, Fennoscandia,
the British Isles and Siberia e.g. /Ekman 1991/; areas where this process is still active today, some
10–15 kyrs after the last deglaciation. In previously glaciated terrain without strong tectonism, glacialisostatic adjustment is the most significant geodynamic process governing vertical deformation of
the crust, e.g. /Peltier 1994/.
During the Last Glacial Maximum (LGM), around 20 kyrs ago, the Laurentide ice sheet of Canada
and the United States, and the Fennoscandian ice sheet had maximum thicknesses of around 2.5–3 km.
When these large ice masses slowly formed, their weight resulted in a slow downwarping of Earth’s
crust. One important factor that governed both this process and the following glacial unloading is thephysical properties of Earth’s crust and mantle. In the downwarping process, mantle material has to be
displaced and flow laterally in order to make room for the flexing crust. At times of ice sheet decay,
mantle flow is reversed and the crust rebounds. Since the mantle viscosity is high, the downwarpingand subsequent rebound are slow processes. Furthermore, it has been shown that a deglaciation of large
Northern Hemisphere ice sheets results in deformation of the Earth’s entire surface, producing a seriesof upwarps and downwarps away from the areas of the former ice sheets (Figure 3-24). However,the deformation is largest in the regions that were glaciated.

Figure 3-24. Calculated deformation resulting from an instantaneous unloading of Northern Hemisphere lastglacial cycle ice sheets at 16 kyrs BP. The results show five sea-level zones with similar sea-level histories
within each zone. Sea-level zone 1 is marked in yellow. The entire Earth is affected by the unloading througha series of elevated and subsided regions away from the areas of the former ice sheets. Modified from /Clark
et al. 1978/.

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The glacial isostatic adjustment process manifests itself not only in the slow rebound regions of past
ice sheets. Current melting of glaciers produces additional GIA effects. As an example, Iceland iscurrently undergoing rapid glacial rebound due to a mass loss of Vatnajökull and other smaller ice
caps e.g. /Árnadóttir et al. 2009/. Since the viscosity of the mantle is inferred to be three orders ofmagnitude lower beneath Iceland /Árnadóttir et al. 2009/, than below for example Fennoscandia,
uplift rates on Iceland are on the order of 20 mm/year, in spite of the much smaller volume of ice loss.
Outside the ice sheet margin, an uplifted forebulge, or peripheral bulge, is formed e.g. /Mörner 1977,
Fjeldskaar 1994, Lambeck 1995/. The forebulge is caused by flexure and a lateral displacement ofmantle material extending outside the ice margin, and it may stretch for several hundreds of kilometres
beyond a major ice sheet. The uplift of the forebulge is considerably smaller than the downwarpingof the crust beneath the central parts of the ice sheet; on the order of tens of metres. During and after
deglaciation, the area of the forebulge experiences land subsidence, exemplified by the ongoing
lowering of the Netherlands, southern England and the east coast of the United States. It is worth
noting that land subsidence in these regions is also a result of ongoing marine sediment loading. The
location of maximum forebulge uplift migrates toward the formerly glaciated region as the ice sheet
withdraws. In addition to the formation of the forebulge, the elasticity of the lithosphere may resultin a downwarping of the crust, not only under the ice sheet, but also to some extent outside the ice
margin. This produces a flexural depression between the ice margin and the forebulge, a depression
where lakes may form from glacial meltwater.
At present, the crust beneath the Antarctic and Greenland ice sheets is depressed in a similar way as
previously occurred under the Laurentide, Fennoscandian and Siberian ice sheets. In Greenland andAntarctica, the crust would also be subject to significant glacial rebound if these areas were to be
deglaciated in the future.
The geoid is the equipotential surface of the Earth’s gravitational field which best fits the undisturbed
surface of the oceans. Shoreline migration is the result of changes in the height of the geoid relative
to the solid surface of the Earth. Sea-level is defined to be zero on land and positive in the oceans,
where it corresponds to the depth of the ocean. The surface of the ocean corresponds to the geoid
surface (see Figure 3-25). Relative sea-level is the vertical height difference between the geoid at anarbitrary time and the present height of the geoid, where the height of the geoid is measured relative
to the height of the solid surface (see Figure 3-25). Relative sea-level is defined to be zero at thepresent day. It is positive during transgression, when the intersection of the geoid with the solid
surface is higher than present, and negative during regression, when the intersection of the geoidwith the solid surface is lower than present. The shape of the geoid is governed by the gravitational
field of the Earth and varies over time. Over time scales of 100 years to 100 kyrs, changes in relative
sea-level throughout Fennoscandia are dominated by the process of glacial isostatic adjustment.
(a)

(b)

t1

SL=0
SL>0

geoid/ocean
surface t0

at t1: RSL>0t0
at t2: RSL<0

Present day (t0):RSL=0 everywheret2

Figure 3-25. (a) At time t0 sea-level (SL) is zero where the depth of the sea is zero, and positive where thedepth of the sea is positive. Sea-level is defined to be zero on land. (b) At time t1 and t2 relative sea-level(RSL) is higher and lower respectively, as a result of the processes described in the text. Note that thechange in relative sea-level is not uniform over the oceans, due to changes in the shape of the geoid.

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Relative sea-level changes associated with GIA are a result of the coupled phenomena isostasy andeustasy. Isostasy is the response of the solid Earth to loading or unloading by ice or water, and/or
unloading and loading due to denudation and sedimentation, whereas eustasy refers to changes in
global sea-level arising from a change in ocean water volume following a mass exchange between
continental ice masses and the oceans.
Global relative sea-level changes exhibit complex spatial and temporal patterns, and are strongly
dependent upon the location relative to major ice sheets. Relative sea-level change at locations far fromice sheets (hereafter referred to as far-field locations) is dominated by the eustatic signal; during deglaciation monotonic sea-level rise causes land inundation and the shift of shorelines inland. At locations
close to ice sheets (hereafter referred to as near-field locations) the isostatic signal dominates; reboundof the solid surface from the time of deglaciation onwards causes land emergence and the migration ofshorelines oceanwards. These are two end-member cases, and, in general, sea-level change, and hence
shoreline migration, is governed by a complex interplay of isostatic and eustatic processes, operating
on different time scales.
GIA-induced sea-level changes depend on the following factors (see Figure 3-26):• The location and thickness of ice sheets.
• The depth and extent of the oceans.
• The structure and properties of the solid Earth and its response to surface loading.
Changes in surface loading arise due to the exchange of mass between ice sheets and ocean basinsthroughout a glacial cycle. The presence of ice-dammed lakes and the redistribution of sediments
also contribute to changes in surface loading, although, in Fennoscandia, the short-lived nature of
ice-dammed lakes, and their shallow depth, means that this perturbation to the pattern of surface
loading probably has a negligible effect when considering the solid Earth response. The denudationof bedrock and redistribution of associated sediments persists over a much longer time scale, on theorder of millions of years. On a regional scale, over a single glacial cycle, the effect of sedimentredistribution upon solid Earth deformation is negligible in relation to the ice load, but over several
glacial cycles, or locally, it may be of significance.
The Earth consists of the crust, the upper and lower mantle, and the outer and inner core. The lithosphere comprises the crust and part of the upper mantle. The average thickness of the lithosphereis c. 100 km, although this value varies between less than 30 km for oceanic lithosphere and up to
more than 200 km for continental lithosphere /Watts 2001/. Lithospheric thicknesses in Fennoscandia
range between ~60 km in the north-west and ~200 km in the south-east /Watts 2001/. The lithosphereis less dense than the mantle below and responds approximately elastically to forces applied at the
GLACIAL ISOSTATIC ADJUSTMENT
Surface Mass
Redistribution• location and thickness of ice sheets• depth and extent of oceans

Earth Structure• Relative sea level

Earth Response• Geopotential• Rotation vector• 3D solid surface
deformation

Figure 3-26. The theory of Glacial Isostatic Adjustment (GIA): inputs and outputs of the system.

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surface. The deformation of this layer is short-lived once the load is removed. The mantle responds
viscoelastically to forces applied over a time scale of ~100 kyrs. The recovery of this layer to a stateof isostatic equilibrium following loading takes several orders of magnitude longer than the recoveryof the lithosphere. The rheological properties of the lithosphere and upper and lower mantle determine
the precise magnitude and duration of solid Earth deformation during a glacial cycle, and are discussed
further in Sections 3.3.2 and 3.3.4. The outer and inner cores are not affected by GIA processes.
GIA-induced sea-level changes arise as a result of the gravitationally-consistent redistribution ofwater between ice sheets and ocean basins, and thus the evolution of surface loading. Any redistribution of surface mass alters the shape of the geoid, which in turn defines the position of the surface ofthe oceans.
The height of the geoid, or mean ocean surface, is dependent upon direct and indirect geoid perturbations /Milne et al. 2002/, as well as changes in ocean water volume and capacity of ocean basins. Direct
effects refer to the deflection of the geoid due to the direct attraction of surface mass loads, such as ice
sheets. Indirect effects refer to geoid perturbations arising due to the surface load-induced deformation.
The volume of the oceans will vary as water is transferred to and from the ice sheets. Also changes inocean capacity and bathymetry arise as a result of crustal and geoidal perturbations in response to ice
and ocean loading. The ice-loading solid surface response includes the depression and rebound of thesolid surface in the location of ice sheets, and the raising and lowering of glacial forebulges in areas
surrounding the ice sheets. The ocean-loading response is similar to the ice-loading response. However,at continental margins ocean loading induces a levering of the continental lithosphere and a subsidence
of offshore regions, this is referred to as continental levering /Clark et al. 1978/ (see Figure 3-27).
The combination of ice and ocean loading leads to a decrease in the volumetric capacity of the ocean
basins during glaciation, and an increase during deglaciation due to the growth and decay of offshore
peripheral bulges, resulting in a globally uniform rise or fall in sea-level, respectively. The ongoing
fall in sea-level following the last deglaciation due to the migration of water from equatorial regions to
subsiding peripheral bulge regions has been termed equatorial ocean syphoning /Mitrovica and Peltier
1991, Mitrovica and Milne 2002/.
Changes in relative sea-level result in shoreline migration. A rise in relative sea-level and the consequentinland migration of the shore may be due to solid Earth subsidence, an increase in the height of the geoid/ocean surface as a result of ice sheet melting, or a combination of these processes. Similarly, a fall in
relative sea-level and the consequent migration of the shore towards the ocean may be due to isostatic
rebound, a fall in the height of the geoid/ocean surface due to ice sheet build up, or a combination ofthese processes. In the vicinity of the Fennoscandian ice sheet, the effects of isostatic deformation anda)Continent

Ocean

t1

b)

Upwarpingt2 > t1

Subsidence

Figure 3-27. Continental levering: the migration of water into offshore regions following deglaciation
results in offshore subsidence and onshore upwarping. Ocean floor subsidence results in sea-level fall at
far-field sites.

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changes to the height of the geoid/ocean surface occur simultaneously, and may have opposite effects
on the position of the shoreline. Further, the surface of the Baltic Sea does not solely depend on theheight of the geoid/ocean surface, but also on the elevation of its sills relative to the Atlantic sea-level,
and ice thicknesses at locations with potential connections to the sea; these may cause it to become
isolated from the global ocean.
Indirectly, groundwater flow and composition are influenced by shoreline migration. If the site is not
subject to permafrost, and it is not covered by an ice sheet, the sea, or a lake, then the groundwater
surface follows a subdued version of the topography. Groundwater flow will be driven by topography,
assuming precipitation exceeds evaporation. If the site is submerged the situation will be almost
stagnant and groundwater flow will be driven by density only /SKB 2010c/. Groundwater composition
is affected, since relative sea-level affects the salinity of the Baltic Sea. Relative sea-level and theextent of the ice sheet determine whether there is a connection between the Baltic Sea and the ocean.
This together with the runoff to the Baltic basin, determines the salinity in the sea/lake water.

3.3.2

Controlling conditions and factors

Loading

The main factors governing the evolution of relative sea-level, and hence shoreline migration, throughout
Fennoscandia are the loading history – both with regard to ice and ocean water – and the rheologicalparameters governing the response of the solid Earth to such a load.
Near-field relative sea-levels are very sensitive to variations in the evolution of the near-field ice sheet,/Lambeck et al. 1998, Davis et al. 1999, Tamisiea et al. 2001, Milne et al. 2002, Kaufman and Lambeck
2002, Tamisiea et al. 2003/. There remain uncertainties in the details of ice loading during the lastglacial cycle, especially with regard to estimates of ice thickness within Fennoscandia.
The details of far-field ice sheets are irrelevant for the prediction of relative sea-level change inFennoscandia /Peltier 1998, Mitrovica et al. 2001b, Tamisiea et al. 2003, Bassett et al. 2005/; masschanges in the far field mainly provide an influence on global eustatic sea-level change. However, theoverall characteristics of far-field ice sheets may also generate a long wavelength isostatic response
in Fennoscandia. This response is likely to be dominated by the signal from a massive North American(Laurentide) ice complex. Due to the position of Fennoscandia on the Laurentide forebulge during theLGM /Mitrovica et al. 1994/, the presence of the Laurentide ice sheet is predicted to have generated
25–30 m of solid Earth uplift in central Fennoscandia at this time (see Section 3.3.4). The magnitude ofthis signal does not vary greatly across Fennoscandia (unlike the relative sea-level signal due to local
ice loading), and is not dependent upon the local geometry of the Laurentide ice sheet; only its volume.
During the last deglaciation an ice-dammed lake prevailed within the Baltic depression when non icecovered sills were above sea-level and all other potential connections between the Baltic Sea andthe North Sea were dammed by the ice sheet ice e.g. /Björck 1995, Lambeck 1999/. The surface of theice-dammed lake was constrained by the height of the ice sheet above sea-level and the surrounding
topography, i.e. the sill levels, and there is evidence to suggest that immediately prior to the drainingof the so called Baltic Ice Lake its surface was 25 m above the contemporary relative sea-level /Björck
1995/. Due to the relatively small volume of water released to the oceans as the ice lake drained theresulting perturbation of global sea-levels was insignificant /Lambeck 1999/. The shallow depth andthe short duration of the ice lake mean that its impact on isostasy was negligible in comparison with
the ice load.
Another process that may impact the loading is erosion and sediment transfer. Erosion and transportof sediment took place beneath warm-based parts of the ice sheet during previous glacial cycles, andthis mass redistribution affects isostatic loading and topography. However, the average total depth ofglacial erosion over all Late Pleistocene glacial cycles in lowland Precambrian parts of Fennoscandia ison the order of a few tens of metres, see Section 3.5. Therefore, on the time scale of one glacial cycle,
the average unloading/loading effect of this process may be neglected. However, any local reshaping of
the land will affect the evolution of the shoreline, and this needs to be taken into account.

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Rheological parameters and topography

In order to determine the Earth’s response to surface loading its internal rheological and density structure
must be specified. The characteristic time scale for loading during a glacial cycle will excite both
elastic and viscous responses, therefore a viscoelastic (Maxwell) rheology is adopted. Such a systembehaves elastically on a short time scale and viscously on a long time scale when placed under stress.
Parameters describing the rheological properties of the Earth’s lithosphere and upper and lower mantle
define its response to loading during a glacial cycle. The average density and elastic structure of
the Earth are taken from /Dziewonski and Anderson 1981/. The lithosphere is generally regarded aselastic for the purposes of GIA studies; this has been shown to be a good approximation throughout
the development of the subject /McConnell 1968/.
The upper mantle lies below the lithosphere, extending to a depth of 660 km, and then the lower
mantle extends to 2,900 km below the Earth’s surface. The viscosity of the upper and lower mantle
have been constrained to lie in the ranges 1·1020–2.6·1021 Pa s and 2·1021–1·1023 Pa s, respectively, as
determined from previous GIA studies /Davis and Mitrovica 1996, Mitrovica and Forte 1997, Simonsand Hager 1997, Lambeck et al. 1998, Davis et al. 1999, Milne et al. 2001, 2002, 2004, Mitrovica andForte 2004/. Lateral variations in radial viscosity structure may be derived from seismic shear-velocity
models by converting velocities to temperatures, and then using temperature to estimate viscosity.
Present-day topography is used to constrain palaeotopography. When calculating palaeotopography,
see Section 3.5, we assume that all changes to the shape of the land arise from the differential GIA
response across the region; however topography is also affected by tectonic, erosional and depositional
processes.
The isostatic adjustment and shoreline migration described above may affect a number of geospherevariables of importance for a deep geological repository (Table 3-3).

3.3.3 Natural analogues/observations in nature
The GIA process may be observed by studying relative sea-level markers, both from the geologicalrecord and from tide gauge data, GPS observations of the 3D deformation of the solid surface, thetime-variation of the gravity field as observed by satellites and land-based gravity surveys, and changes
in the orientation of the Earth’s rotation vector and length of day.
Geological and tide gauge data complement each other. Tide gauge data consist of monthly or annual
observations of mean sea-level relative to a local solid Earth marker at coastal sites throughout
the world. The data may be accessed through e.g. the “Permanent Service for Mean Sea-level”
at www.pol.ac.uk/psmsl. In the Baltic region, such sea-level observations have been conductedsystematically for more than a century. When measuring isostatic uplift rates with this method, it isnecessary to correct the result in order to exclude the effect from eustatic variations. In non-coastal
areas, present post-glacial rebound rates have been estimated by repeated geodetic levelling, tied tolocations of sea-level observation. The results from these measurements show that the present uplift
rate displays a concentric uplift pattern over Fennoscandia, with a maximum rate of slightly morethan 9 mm/yr in the northernmost part of the Baltic Sea, see Figure 3-28.
The geological data, including palaeoshoreline positions, lake isolation and tilting information,
cover a longer time period than the tide gauge data, but the tide gauge data are more accurate, bothin terms of their vertical resolution and the dating of the information. In this context, shorelines
or other geomorphological features that formed at the highest post-glacial sea-level, in front of theretreating ice sheet margin, provide the oldest estimate of former relative sea-level positions.
However, the total amount of post-glacial uplift at a site is typically larger than can be inferredfrom, for example, raised beaches. A significant portion of the uplift takes place as the ice sheet
starts to decay, prior to the actual deglaciation of a typical site situated at some distance from the
maximum ice margin. The total maximum amount of glacial rebound that has occurred due to the
decay of the Fennoscandian ice sheet is around 800 m, e.g. /Mörner 1979/ (Figure 3-28). Thismay be compared to the largest value on rebound as inferred from the highest marine limit, which
is situated at ~280 m a.s.l. in the Swedish coastal region of the Gulf of Bothnia. Another related
method to study glacial rebound is to analyze the amount and direction of tilt of palaeoshorelinesof glacial lakes that formed behind the retreating ice margin. A selection of geological relative80

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Table 3-3. A summary of how geosphere variables are influenced by isostatic adjustment andshoreline migration.
Geosphere variable

Climate issue variable

Summary of influence

Rock stresses

Isostatic depression/reboundThe deformation of the Earth’s crust will lead to altered rockstresses.

Groundwater pressure

Relative sea-level

If the site is not covered by the sea or a lake the groundwaterpressures will be determined by topography and groundwater
recharge. If the site is submerged the groundwater pressure
will be determined by depth of the sea/lake.

0250

500 km

800
700
600
500
400
300
200
100
0

-100

Figure 3-28. Total amount of glacial rebound (metres) that has taken place due to the decay of the
Fennoscandian ice sheet. The bold line shows the approximate maximum extent of the last Fennoscandian
ice sheet. Note that the total amount of rebound is more than twice as large as the amount of reboundthat may be inferred from raised beaches formed at the highest post-glacial sea-level (see the text). After/Fredén 2002/.

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sea-level data for Fennoscandia can be found in e.g. /Lambeck et al. 1998, Påsse 2001, Eronen et al.
2001, Kaufmann and Lambeck 2002, Whitehouse 20072/, and references therein, but at present acomplete compilation of such data does not exist.
Care must be exercised when interpreting palaeoshoreline data from the Baltic Sea; it is important todetermine whether the data relate to sea-level or a lake level. An overriding problem with geological
data in Fennoscandia is the lack of relative sea-level data prior to the LGM. Most of the evidence ofshorelines prior to this time has been destroyed by the ice sheet. This makes it difficult to test relative
sea-level predictions from before 20 kyr before present. Furthermore, the fact that shorelines canonly form in ice-free locations provides an important temporal constraint for the testing of shorelinemigration predictions against undated shoreline data.
As previously mentioned, observations of relative sea-level are required in both the near and far
field. Far-field data provide constraints on eustatic sea-level change and hence the volume of ice
contained in far-field ice sheets. Near-field data exhibit a more complex dependence upon the local
distribution of ice and the Earth model.
One way of making direct observations of ongoing post-glacial crustal deformation is to use highquality data from networks of continuously operating permanent GPS receivers. GPS data provide
satellite-measured observations of changes in baseline distances which yield present-day rates of verticaland horizontal motion at a series of discrete positions. GPS data from the BIFROST project /Johansson
et al. 2002/ provide excellent spatial coverage of present-day solid Earth deformation throughout
Fennoscandia to a high degree of accuracy. Typically, the GPS stations used in the BIFROST project
were established within national land survey programmes with an initial aim of providing reference
coordinates for other GPS measurements. The detailed analyses of data from such permanent GPS
stations have provided new insight into the processes of post-glacial rebound or GIA in Fennoscandia
and Canada, cf. /Scherneck et al. 2001, Henton et al. 2006/. The results provide information on both
the rate of the vertical uplift component, as well as on the associated smaller horizontal component ofcrustal motion, e.g. /Johansson et al. 2002/ (Figure 3-29). The rates are averaged over the period ofdata collection, which is generally on the order of 10 years at the present, to yield estimates accurate
to within ~0.1 mm/yr in the horizontal direction, and a few tens of a mm/yr in the vertical direction.
Uplift determined by these GPS observations show the same concentric uplift patterns as thosederived from sea-level and levelling measurements. The fastest rebound occurs approximately inthe areas where the Laurentide and Fennoscandian ice sheets had their greatest thicknesses. The
maximum vertical uplift rate measured in this way in the area of the former Laurentide ice sheet
is ~13 mm/yr /Henton et al. 2006/, while the corresponding value for Fennoscandia is ~11 mm/yr(Figure 3-29). The largest horizontal displacements are generally found in the area of the ice marginsof the maximum extent of the ice sheet. The difference between the maximum vertical uplift rates
as observed from sea-level/levelling measurements and from the analysis of GPS data is to a large
extent covered by the uncertainty errors of the measurements, mainly in the sea-level measurements.
Since the GPS network yields an estimate for the full 3D deformation field, and the horizontal andvertical components of this field have different sensitivities to the ice history and Earth model, thisopens up the possibility of constraining model parameters more accurately when these data are
combined with sea-level observations. Combining the GPS data with sea-level data is particularly
important when one considers that the GPS data are limited by their relatively short time span, andcan only yield information about present-day deformation rates. Also, sea-level data are limitedby their poor spatial coverage; a factor that is addressed by the distribution of GPS observations
throughout the interior of Fennoscandia.
Another, indirect, way of studying ongoing post-glacial rebound is by absolute gravity measurements,
e.g. /Lambert et al. 2001, Mäkinen et al. 2005/, a method that may be used to estimate also the remaining
uplift in areas where the process is not complete.

Whitehouse P, 2007. A relative sea-level data base for Fennoscandia. Durham university, February 2007.
Excel data base stored at SKB, SKBdoc 1265613.

2

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Vertical deformation rates (mm/yr)-2 0 2 4 6 8 10 12 14

70°.

Horizontal rates, error ellipses,and vertical error bars (mm/yr)

1mm/yr(1σ)

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5

7

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11

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60°

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20°

30°

Figure 3-29. Present-day crustal deformation over Fennoscandia as observed by continuous GPS measurements within the BIFROST project, e.g. /Scherneck et al. 2001/. The rebound displays a concentric patternwith a maximum rebound rate of ~11 mm/yr located approximately in the area of the former maximum ice
sheet thickness. In the post-glacial upwarping process, a horizontal component of crustal deformation is alsopresent (blue arrows), directed outward from the area of maximum depression. Modified from /Lidberg 2007/.

3.3.4

Model studies

The GIA model

For a comprehensive and detailed review of present day GIA models, see /Whitehouse 2009, Chapter 4/.
That chapter includes descriptions of different GIA approaches used by various research groups andrelative accuracy of these methods. Recent improvements of the GIA theory are described, as well as
current shortcomings of the models. The various data sets used to calibrate and verify the accuracyof the modelling are also briefly discussed. A shorter description of GIA models are given here,
followed by the GIA simulations performed for the safety assessment.
The majority of GIA models solve the sea-level equation, which was originally developed by /Farrell
and Clark 1976/, and describes the gravitationally-consistent redistribution of water from ice sheets

TR-10-4983to ocean basins, and thus the evolution of the distribution of water between oceans and land-based
ice sheets. Any redistribution of surface mass alters the shape of the geoid, which in turn defines theredistribution of water in the ocean basins; therefore an iterative procedure is required to solve thesea-level equation.
There have been many attempts to construct a dynamical model to explain the observed processesof GIA. Following /Farrell and Clark 1976/, early studies neglected the time dependence of ocean
shorelines /Wu and Peltier 1983, Nakada and Lambeck 1989, Tushingham and Peltier 1991/, and theice loading was simply applied as a series of ‘finite elements’ /Clark et al. 1978, Peltier et al. 1978,
Wu and Peltier 1983/. In 1991, /Mitrovica and Peltier 1991/ derived a pseudospectral approach tosolving the sea-level equation, and this method has been widely used throughout the field ever since.
Models which take into account the time-varying nature of shorelines began to be developed in theearly 1990s /Lambeck and Nakada 1990, Johnston 1993, Peltier 1994, 1998, Milne 1998, Milne et al.
1999, Peltier and Drummond 2002, Kendall et al. 2005/, and the effects of Earth-rotation have also
subsequently been considered /Han and Wahr 1989, Bills and James 1996, Milne and Mitrovica 1998,
Peltier 1998, Mitrovica et al. 2001a, Mitrovica et al. 2005/. Initial investigations into the effect oflateral variations in Earth structure are also in the process of being developed /Kaufmann et al. 2000,
Wu et al. 2005, Latychev et al. 2005b, Paulson et al. 2005, Whitehouse et al. 2006/.
The GIA model used in this study was developed by Milne /Milne 1998, Milne and Mitrovica 1998,
Milne et al. 1999/. Three refinements to the original sea-level equation presented by /Farrell and
Clark 1976/ have been developed by /Mitrovica and Milne 2003/. Firstly, time-dependent shoreline
positions are taken into account when calculating the ocean-loading function. Shorelines can change
position by several hundreds of kilometres in flat terrains, and this must be accounted for in applying
the ocean load. Secondly, the water influx into regions vacated by retreating, marine-based ice iscarefully accounted for in the distribution of the load /Milne et al. 1999/. And thirdly, changes to therotational state of the Earth as a result of both surface and internal mass redistributions are considered.
The model used to generate predictions of relative sea-level change in this study is based on a newgeneral theoretical foundation that includes all these advances.
The sea-level equation

The sea-level equation takes account of changes to the height of the geoid and the Earth’s solid surface.
Local factors, such as changes to the tidal regime, the consolidation of sediments, and tectonic processes,
are neglected in this study.
Δξrsl(τ,φ) = Δξeus(τ)+Δξisos(τ,φ)+Δξlocal(τ,φ)

Equation 3-9

In the above expression the left-hand side refers to changes in relative sea-level at time τ and location φ.
The first term on the right-hand side is the time-dependent eustatic signal, the second term varies inspace and time, and relates to the isostatic effects of glacial rebound, including both ice and waterload contributions, and the third term refers to local factors, as described above. In order to solve
Equation 3-9 in a gravitationally self-consistent manner, a pseudospectral algorithm /Mitrovica andPeltier 1991, Milne and Mitrovica 1998/ is employed. Green’s functions are constructed to determine
the GIA-induced perturbations to the geopotential and solid surfaces due to loading. The resulting
temporal convolutions are evaluated by describing the GIA loading history as a series of discrete
Heaviside increments. The spatial convolutions are performed by transforming the problem to thespectral domain and employing the pseudospectral algorithm /Mitrovica and Peltier 1991, Milneand Mitrovica 1998/.
The method by which relative sea-level is calculated at a certain position, at a certain time, may be
broken down into a series of simplified steps. At the start of the model run, loading is applied to an
Earth model that is assumed to be in isostatic equilibrium. Once an ice load is applied at each time
step the resulting deformation of the solid Earth and the perturbation to the geoid are calculated. The
new shape of the geoid determines the redistribution of water in the oceans and the new extent of theoceans. However, this redistribution of water in turn affects the shape of the geoid and the deformation of the solid Earth in oceanic regions, therefore an iterative procedure is used to ensure the correct
treatment of perturbations to the solid and geoid surfaces; recalculations of the perturbation to thegeoid and the redistribution of water are carried out until there are no further changes, at which point
loading for the next time step is applied. Loading is applied via forward time stepping.
84

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Earth model
Sea level
equation

Sea level curves;
relative to present
day sea level

Ice model

Topography
data

Shoreline positions at
discrete times definedby zero palaeotopography
contour

Topography
calculation

Palaeotopography
at discrete times

Figure 3-30. Flow chart outlining the inputs (in red), calculations (in green) and outputs (in blue) required
to calculate palaeoshoreline positions due to GIA processes.

The final output of the GIA model consists of a series of global relative sea-level predictions for
all time steps used in the model. In the method used here, relative sea-level is defined to be zero at
the present day, and heights are given relative to this ‘zero level’ at all times in the past and future.
Corrections are applied using present-day topography to determine the height of any point above or
below sea-level at any given time. Assuming that the underlying topography is unaltered by erosion
and sedimentation the evolution of palaeotopography is calculated via the following equation:T(φ,τp) = T(φ,τ0) – ξrsl(φ,τp)						Equation 3-10
where T is topography, defined to be the height of the solid surface above sea-level at time τ, andξrsl is relative sea-level at time τ. τp refers to a time in the past, τ0 refers to the present day, and φ isthe location on the surface of the Earth. Shoreline positions at time τp are determined by the zero
palaeotopography contour at that time.
Ice loading

The global ice loading model used in this study is modified from the ICE3G deglaciation history/Tushingham and Peltier 1991/, and has been calibrated using far-field relative sea-level data /Radtke
et al. 1988, Fairbanks 1989, Bard et al. 1990, 1996, Chappell and Polach 1991, Chappell et al. 1996,
Hanebuth et al. 2000, Yokoyama et al. 2000/. A eustatic curve has been used to tune the mass of ice
contained within far-field ice sheets. The near-field ice loading for Fennoscandia is the one describedin Section 3.1.4 and has been derived using a thermodynamic ice sheet model employing a proxy
data palaeotemperatures curve. From this model, the extent and thickness of the Fennoscandian ice
sheet at a series of discrete times from 116 kyr BP to the present day has been derived and used.
Ice loading history within Fennoscandia is the principal factor governing relative sea-level change inFennoscandia. The loading history presented in Section 3.1.4 is perturbed to investigate the sensitivityof sea-level change to differences in ice thickness, the timing of deglaciation, the pattern of ice build-up,
and the time scale over which loading is considered (Table 3-4). The extent of ice at each time stepsince the LGM has not been perturbed because the geometry of the ice sheets during this period are
relatively well constrained by geological data. Details of the ice build-up prior to the LGM is less wellconstrained because geological evidence subsequently has been destroyed. However, for consistency,
the ice extents for this period are also not altered. An Earth model consisting of a 96 km-thick elastic
lithosphere, an upper mantle of viscosity 0.5·1021 Pa s, and a lower mantle of viscosity 1·1022 Pa s isused to investigate the response to the different loading models.

TR-10-49

85

Table 3-4. Summary of ice loading models used in this study. In all cases the ICE3G global ice
model is used to constrain the distribution and thickness of ice outside Fennoscandia.
Model

Description of loading model

Model run time1

Weichselian glacial cycle as described in Section 3.1.4

116 kyr2

Two Weichselian glacial cycles run after each other to give a double glacial cycle

232 kyr3

Weichselian glacial cycle, loading only applied from 40 kyr BP

40 kyr4

Weichselian glacial cycle with 90% ice thickness in Fennoscandia

116 kyr5

Weichselian glacial cycle with 80% ice thickness in Fennoscandia

116 kyr6

Weichselian glacial cycle with the timing of deglaciation advanced by 500 years

116 kyr7

Weichselian glacial cycle modified with ice thickness allowed to increase linearly
between 60 kyr BP and 20 kyr BP

116 kyr

8

Four relaxation models, as described in the isostatic memory section

Varies9

Weichselian glacial cycle with all ice removed from Fennoscandia for the duration of themodel run, as described in the section on far-field ice sheets

116 kyr

In Figure 3-31a the same Fennoscandian surface loading is used, but the model run is initiated at
different times. Due to the assumption that the Earth is initially in isostatic equilibrium there will besome discrepancy between relative sea-level predictions at early times in a model started at 116 kyr
BP and one that has already been running for 100 kyr because the Earth’s response is dependent uponloading history as well as the instantaneous load (see the section on Isostatic Memory). Because theEarth has undergone a series of glacial cycles, a loading model that accounts for loading and unloadingof the Earth during the previous cycle will give more realistic predictions for relative sea-level during
the cycle of interest. However, the negligible difference between relative sea-level predictions fromthe single cycle, double cycle and 40 kyr models (models 1, 2, and 3 in Table 3-4), for times between
25 kyr BP and the present day, imply that no long-term error is introduced when a model with a later
start time is used. It is important that the shortest possible time steps are used to ensure that the fullloading history is captured; failure to do so will miss short time scale fluctuations in ice distribution,
leading to a decrease in the accuracy of the relative sea-level predictions.
The effect of altering the thickness of ice in the loading model is illustrated in Figure 3-31b. There islittle difference in the shape and magnitude of predicted relative sea-level curves for the 80%, 90%
and 100% loading models during the minor glaciation between ~65 kyr BP and ~50 kyr BP (models 5,
4 and 1 in Table 3-4), but comparing results from the 80% and 100% models at the LGM yields
differences of ~100 m at Forsmark. The larger the range of ice thicknesses used, the greater the rangeof relative sea-level predictions, as illustrated by the range of curves for Forsmark.
Relative Sea Level - Comparison of ice loading models (Forsmark)
600

(a)

400

200

rsl (m)rsl (m)

400

0
-200

400

200
0
-200

-200 -150 -100

-5 0time (kyr BP)

0

-600

(c)

200
0

-200

-400

-400
-600

600

(b)
rsl (m)

600

-400
-100

-7 5

-5 0

-25

time (kyr BP)

0

-600

-100

-7 5

-50

-25

time (kyr BP)

0

Figure 3-31. Predicted relative sea-level curves for Forsmark. See Table 3-4 for details of the ice loading
models. An Earth model with a thin lithosphere and low upper mantle viscosity has been used in all cases(Model A, Table 3-5). a) black: model 1, red: model 2, green: model 3. b) black: model 1, red: model 4,green: model 5. c) black: model 1, red: model 6, green: model 7.

86

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The timing of maximum relative sea-level is the same for all loading models except that where
the timing of deglaciation is brought forwards by 500 years (model 6 in Table 3-4). In this case,maximum relative sea-level also occurs 500 years earlier. Apart from this time shift, which decays tozero by the present day, there is otherwise no difference between predictions for this model, and theWeichselian glacial cycle. The magnitude of the maximum relative sea-level is identical to that predicted by the single cycle, double cycle and 40 kyr loading models (models 1, 2 and 3 in Table 3-4).
In model 7 (Table 3-4), ice thickness is allowed to increase linearly between 60 kyr BP and 20 kyr
BP. This model has been developed to test the postulation that there was a continuous ice sheet presence in Fennoscandia during this period /Lundqvist 1992/, see also Section 4.2 and 4.4. The patternof continuous, linear ice sheet growth is purely an assumption of this model sensitivity test; themaximum ice thickness of the Weichselian glacial cycle is never exceeded, but the cumulative effectof loading the Earth for a longer period result in a greater amount of isostatic depression throughout
Fennoscandia, and a greater relative sea-level maximum (Figure 3-31c). Therefore the GIA signal is
not only a function of the thickness and timing of ice loading at a given time, but also the evolutionof the loading.
The rate of change of relative sea-level throughout the glacial cycle is similar for all the loading models
except the model where the pattern of ice sheet growth has been markedly altered prior to the LGM(model 7 in Table 3-4, Figure 3-31c). However, predictions of present-day uplift rates for the various
models do highlight small differences between the models (see Figure 3-32). Uplift rates for the single
cycle, double cycle, 40 kyr loading model and the model where the timing of deglaciation is shiftedby a small amount are virtually identical. Much lower rates are predicted for the case where only
90% or 80% of the ice thickness is used, although the lateral extent over which rebound occurs issimilar to the standard loading models. The final model, where the pattern of ice sheet growth has
been altered, yields much greater present-day uplift rates, reflecting the fact that the solid Earth has
to rebound from a position of greater isostatic depression at the LGM. In this final case, the area overwhich rebound is still taking place is greater than that predicted by the other loading models.
Comparisons between predictions from the GIA model, using the ice loading from the Weichselian
glacial cycle as outlined in Section 3.1.4 (model 1 in Table 3-4), and observations of relative sea-leveland present-day uplift rates in Fennoscandia show that the modelling overpredicts the GIA response(see Figures 3-33 and 3-34). There may be several reasons for this; either the ice load has been toolarge in the ice load history, as tested above with load models 4 and 5 reducing the Weichselian glacial
cycle ice sheet thickness to 90% and 80% respectively, or the misfit could be a result of assuming
a laterally homogeneous (1D) Earth structure, discussed in the section A GIA case study with a
Fennoscandian 3D Earth structure below, or a combination of the two.
Unlike the geological evidence relating to the spatial extent of the Weichselian ice sheet, very few
constraints upon the ice sheet thickness exist, and one explanation for the misfit could be that ice
thicknesses in the loading model are too large. As pointed out by /Denton and Hughes 1981/, early
ice sheet modelling studies of steady-state ice sheets, including their own, produced ice thicknesses
that were greater than one would expect from ice sheets during a natural glacial cycle. This has since
been confirmed by GIA modelling, showing that those thick ice sheet profiles result in a poor fit
to relative sea-level observations. Subsequent dynamical approaches to ice modelling outlined in
Section 3.1, including the model used in the present study, produce considerably thinner ice sheets,for example at the LGM. These models yield a closer fit to relative sea-level observations when used
as an input to GIA models, but in the present study there is still a misfit to explain. Some of thisdifference may be explained by the fact that there are still gaps in our knowledge of basal processes
related to, for example, sliding and sediment deformation under ice sheets. Another explanation
could be that the assumption of a uniform Earth structure results in a too large GIA response, see
below.
Earth rheology

The second main factor governing sea-level change, following the load history, is the Earth model.
A three-layer, 1D radial structure is used in these experiments. A discussion of the reasonable
joint parameter range for the rheological properties of the lithosphere and mantle can be found in
Section 3.3.2. The details of the models used in this investigation are outlined below in Table 3-5.

TR-10-49

87

Present day uplift rates - Comparison of ice loading models
70˚

(a)

70˚

(b)

70˚

65˚

65˚

65˚

60˚

60˚

60˚

55˚

10˚

20˚

70˚

55˚

30˚

10˚

(d)

20˚

70˚

65˚

65˚

60˚

60˚

55˚

70˚

10˚

20˚

30˚

(f)

55˚

70˚

65˚

65˚

60˚

60˚

55˚

10˚

20˚

30˚

55˚

30˚

55˚

(c)

10˚

20˚

30˚

(e)

10˚

20˚

30˚

20˚

30˚

(g)

10˚

-5 -4 -2 0 2 4 8 1012141618

uplift rate (mm/yr)

Figure 3-32. Predicted present-day uplift rates throughout Fennoscandia for the various ice-loading models.
See Table 3-4 for details of the loading models. An Earth model with a thin lithosphere and a low upper
mantle viscosity has been used in all cases (Model A, Table 3-5). a) Loading model 1. b) Loading model 2.
c) Loading model 3. d) Loading model 4. e) Loading model 5. f) Loading model 6. g) Loading model 7.

88

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Testing relative sea level data against GIA predictionsSmålandAngerman0˚rsl (m)

400
300

5˚

10˚

15˚

20˚

30˚

25˚

500

35˚

70˚

400

rsl (m)

500

70˚

200
100

300
200
100

65˚

0
-15

-10

-5

65˚

0

Angerman0-15

time (kyr BP)

-5

0

55˚

Uppsala60˚

UppsalaSmåland
Blekinge 2

Gotland
Blekinge 1

55˚

Gotland

500

500

400

50˚

300

0˚

5˚

200

10˚

15˚

30˚

25˚

20˚

400

50˚
35˚

rsl (m)rsl (m)-10time (kyr BP)
60˚

100

300
200
100

0
-10

-5

Blekinge10

rsl (m)time (kyr BP)0

Blekinge2

500

500

400

400

300

300

rsl (m)

-15

200
100

-15

-10

-5

0

time (kyr BP)

200
100

0

0
-15

-10

-5

0

-15

time (kyr BP)

-10

-5

0

time (kyr BP)

Figure 3-33. Comparison between predicted relative sea-level (black line) and observations from relative
sea-level markers (red points) at six sites in Fennoscandia (green stars) for the last 15 kyr. Data compiled
from /Lambeck et al. 1998/.
Comparison of predicted and observed
present day uplift rates
observed
predicted
70˚

(a)

70˚

65˚

65˚

60˚

60˚

55˚

(b)

55˚
10˚

20˚

30˚

10˚

20˚

30˚

0 2 4 8 10 12 14 16

uplift rate (mm/yr)

Figure 3-34. Comparison of predicted and observed present-day uplift rates. a) Predicted present-day
uplift rates interpolated at GPS sites (black dots and triangles). Data generated using the Weichselian
glacial cycle loading model (model 1 in Table 3-4) and an Earth model consisting of a 96 km-thick elastic
lithosphere, an upper mantle of viscosity 0.5·1021 Pa s, and a lower mantle of viscosity 1·1022 Pa s (model A in
Table 3-5). b) Observed present-day uplift rates interpolated at GPS sites. Data from the BIFROST project/Johansson et al. 2002/.
TR-10-49

89

Table 3-5. A summary of the five Earth models considered in this study.
Model

Model description

Lithosphericthickness (km)Upper mantle
viscosity (Pa s)

Lower mantle
viscosity (Pa s)

AThin lithosphere96

0.5·1021

1·1022

B

Thin lithosphere, strong lower mantle

96

21

0.5·104·1022C

Thick lithosphere120

0.5·1021

1·1022D

Thick lithosphere, strong lower mantle

120

0.5·1021

3·1022

E

Thick lithosphere, strong upper mantle

120

0.8·1021

1·1022For all of the loading models considered here, for the majority of the glacial cycle, there is less than
~20 m difference in predicted relative sea-level across all five Earth models (see Figure 3-35).
However during the deglaciation of the Fennoscandian ice sheet, at around 14.5 kyr BP, the differencein predictions for the five Earth models increases to ~100 m for Forsmark. At any one time, differencesin geoid height at any fixed point will be negligible if the same loading model is used; any differencesin the relative sea-level are therefore due to variations in the solid Earth response to loading.
The divergence in the predictions is short-lived, implying that the phenomenon is related to the elastic
properties of the lithosphere. In general, an Earth model with a thinner lithosphere will undergo thegreatest isostatic deformation, resulting in the greatest relative sea-level prediction. From the results, italso seems that a model with a weaker lower mantle (e.g. models A and C) will undergo greater deformation than one with a stronger, or higher viscosity, lower mantle, while the model with the strongest
upper mantle viscosity (model E), exhibits the smallest amount of solid Earth deformation.
Since the differences in relative sea-level predictions are short-lived, and very few relative sea-level
data are available from this period, it is not possible to test these predictions against data in order toconstrain the Earth parameters. Instead, present-day uplift rates must be used.
600

400

200

rsl (m)rsl (m)

400

600

(a)

0
-200
-400

(b)

Relative Sea Level Comparison of EarthModels (Forsmark)

200
0
-200
-400

-600

-600
-200 -150 -100 -50-100 -75 -50 -25

0time (kyr BP)

0time (kyr BP)
Key: Earth models

600

400

200

rsl (m)rsl (m)

400

600

(c)

0
-200

A

(d)

B

200

C

0

D

-200

E

-400

-400

-600

-600
-100 -75 -50 -25

time (kyr BP)

0-100 -75 -50 -25

0time (kyr BP)

Figure 3-35. Comparison of Earth models. Predicted relative sea-level curves for Forsmark for all five
Earth models under consideration. See Table 3-4 for details of the loading models. a): Double Weichselian
glacial cycle – loading model 2. b): Single Weichselian glacial cycle – loading model 1. c): Single Weichselian
glacial cycle modified with 80% ice thickness in Fennoscandia – loading model 5. d): Single Weichselian
glacial cycle modified with a constant, linear increase in ice thickness in Fennoscandia between 60 kyr BP
and 20 kyr BP – loading model 7.

90

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Earth models with a thinner lithosphere (models A and B in Table 3-5) are predicted to experiencegreater present-day uplift rates, although the extent over which rebound takes place will be slightly
smaller than for models with a thicker lithosphere (see Figure 3-36 and 3-37). Earth models witha weaker lower mantle (models A and C in Table 3-5) tend to experience greater present-day uplift
rates, over a larger area, than models with the same lithospheric thickness but greater lower mantle
viscosity (models B and D respectively in Table 3-5). The existence of a stronger upper mantle(model E in Table 3-5) also increases the predicted rate of uplift.
In reality, the Earth is not laterally homogeneous as assumed in these simulations, and variations
in lithospheric thickness and mantle viscosity will perturb the GIA signal. /Paulson et al. 2005/
estimated that for realistic variations in mantle viscosity, present-day uplift rates may be altered by
up to 10%. The effect of including a 3D Earth structure for Fennoscandia has been investigated by/Whitehouse 2009, Section 4.5.4/, part of which is presented in the section A GIA case study with a
Fennoscandian 3D Earth structure below.

Present day uplift rates - Comparison of Earth Models
70˚

(a)

70˚

(b)

70˚

65˚

65˚

65˚

60˚

60˚

60˚

55˚

55˚

55˚
10˚

20˚

70˚

10˚

30˚

(c)

20˚

(d)

70˚

65˚

65˚

60˚

60˚

10˚

30˚

20˚

30˚

(e)

55˚

55˚
10˚

20˚

30˚

10˚

20˚

30˚

-5 -4 -2 0 2 4 8 10 12 14 16 18

uplift rate (mm/yr)
Figure 3-36. Comparison of Earth models. Predicted present-day uplift rates throughout Fennoscandia for
the five Earth models under consideration. The loading model is defined to be a single Weichselian glacial
cycle (Model 1) in all cases. a) Earth model A. b) Earth model B. c) Earth model C. d) Earth model D.
e) Earth model E.

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91

Present day uplift rates - Misfit between observations
and predictions
70˚

(a)

70˚

(b)

70˚

65˚

65˚

65˚

60˚

60˚

60˚

55˚

55˚

55˚
10˚

20˚

70˚

10˚

30˚

(c)

20˚

(d)
(d)

65˚

60˚

60˚

20˚

30˚

(e)

70˚

65˚

10˚

30˚

55˚

55˚
10˚

20˚

30˚

-10 -5

-4

10˚

-3

-2

-1

20˚

0

1

30˚

2

uplift rate (mm/yr)
Figure 3-37. Residual present-day uplift rates throughout Fennoscandia for the five Earth models underconsideration. The predictions of Figure 3-36 are interpolated at the BIFROST GPS sites and thensubtracted from the observed signal (see Figure 3-38). Loading and Earth models as in Figure 3-36.

Isostatic memory

The time scale over which the GIA model is run has implications for the accuracy of relative
sea-level predictions due to the viscous memory of the Earth model. The Earth is assumed to bein isostatic equilibrium at the start of a model run. In reality, the Earth is unlikely to reach such astate if the advance and decay of ice sheets continues with the same periodicity as seen in the past
(Figure 1-2). Therefore, the model must be run for a long enough period that sufficient loading history is considered, and the initial condition of isostatic equilibrium does not affect predictions during
the period of interest.

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In order to investigate the recovery of the viscoelastic Earth model following unloading four loading
models were defined.
• The ‘standard’ model: Weichselian glacial cycle loading is applied for a single glacial cycle.
• The ‘double’ model: Weichselian glacial cycle loading is applied over two consecutive glacial cycles.
• The ‘40 kyr’ model: Weichselian glacial cycle loading applied only from 40 kyr BP.
• The ‘80%’ model: Weichselian glacial cycle loading modified with only 80% ice thickness in
Fennoscandia is applied for a single glacial cycle.
In all four cases, the loading models were run using three different Earth models (A, C and D in
Table 3-5 above), and following the final loading time step at 0 kyr BP the relative sea-level calculations
were allowed to run for another 60 kyr into the future without any further change to the distributionof surface loads, thus enabling the system to return to equilibrium.
The vertical solid Earth response was calculated at six sites for each model run. These sites werechosen to lie roughly 250 km apart along a profile running approximately south from the present-day
centre of isostatic uplift (see Figure 3-38) in order to investigate the solid Earth response at a rangeof distances from the previous centre of ice loading, and for a range of loading histories.
Deformation at each of the six sites is defined to be zero at the start of the model run. During theexperiment, solid Earth deformation at each site is measured relative to the initial position, withnegative values indicating a decrease in elevation due to isostatic subsidence. The timing and magnitude of maximum depression is recorded in each case, and when the deflection of the site returns toless than 1% of the maximum deflection, equilibrium is assumed. The isostatic signal at sites 1, 4 and 6
are shown in Figure 3-39, whereas Table 3-6 summarises the results of the model runs.
For a given site and Earth model, there is little variation in the recovery time for the four different
loading models. The shortest loading model is initiated only ~25 kyr before the timing of maximum
deformation, and the difference in solid Earth deformation between this model and the double-cycle
loading model is negligible from the timing of maximum deformation onwards (see Figure 3-39).
Therefore the isostatic response of the Earth models is not dependent upon loading prior to ~25 kyr
before the timing of maximum deformation, and a loading model need only contain information from~25 kyr prior to the period of interest to ensure no distortion of the results due to initial conditions.
(a) Sites used in Isostatic Sensitivity Tests
0˚

5˚

10˚

15˚

20˚

25˚

30˚

(b) Present day uplift rates
from GPS measurements

35˚

70˚

70˚

65˚

70˚

65˚

1
2

65˚

3

60˚

60˚

4
5

55˚

55˚

60˚

6
55˚

50˚
0˚

5˚

10˚

15˚

20˚

25˚

30˚

50˚ 35˚

10˚

20˚

30˚

-4 -3 -2 -1 0 2 4 6 8 1012

uplift rate (mm/yr)

Figure 3-38. a) Map showing the location of the six sites (green stars) where the isostatic response iscalculated and b) the present-day pattern of uplift in Fennoscandia from BIFROST GPS rates /Johansson
et al. 2002/.

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Isostatic Recovery Sensitivity Tests

0
-100
-200
-300
-400
-500
-600
-700

0
-100
-200
-300
-400
-500
-600
-700

50

-200-150-10 0 -50

time (kyr BP)

-300
-400
-500
-600
-700

-100
-200
-300
-400
-500
-600
-700

0
-100
-200
-300
-400
-500
-600
-700

50

-200-150-10 0 -50

point 4, earth model D

0

100
0
-100
-200
-300
-400
-500
-600
-700

50

-200-150-10 0 -5 0 0time (kyr BP)

0
-100
-200
-300
-400
-500
-600
-700
50

point 6, earth model D

100
0
-100
-200
-300
-400
-500
-600
-700
-20 0-150-100 -50

50

time (kyr BP)

point 6, earth model C
isostatic deformation (m)point 6, earth model A
100

50

time (kyr BP)

100

time (kyr BP)

isostatic deformation (m)-200-200-150-10 0 -5 0 0

isostatic deformation (m)0time (kyr BP)

0
-100

50

point 4, earth model C
isostatic deformation (m)isostatic deformation (m)100-200-150-100 -5 0 0

100time (kyr BP)

point 4, earth model A-200-150-100 -5 0 00isostatic deformation (m)-200-150-100 -5 0 0100

isostatic deformation (m)100point 1, earth model D

point 1, earth model C
isostatic deformation (m)isostatic deformation (m)point 1, earth model A

0

50

time (kyr BP)

100
0
-100
-200
-300
-400
-500
-600
-700
-200-150-10 0 -5 0 0

50time (kyr BP)

Key: Loading Models
Double80%

Standard
40 ka

Figure 3-39. Isostatic recovery sensitivity tests. Solid Earth deformation at sites 1, 4 and 6 (see Figure 3-38)
for the four loading models and three Earth models described in the isostatic recovery sensitivity tests.

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Isostatic Recovery Sensitivity Tests - Future Scenarios

-60
-80
-100
-120
0

10 20 30 40 50 60

-20
-40
-60
-80
-100
-120
0

time (kyr BP)

-40
-60
-80
-100
-120
0

10 20 30 40 50 60

-60
-80
-100
-120
0

10 20 30 40 50 60

time (kyr BP)

isostatic deformation (m)isostatic deformation (m)

-40

-60
-80
-100
-120
0

0

-40
-60
-80
-100
-120
0

10 20 30 40 50 60

10 20 30 40 50 60

0
-20
-40
-60
-80
-100
-120
0

time (kyr BP)

point 6, earth model A
0-40

point 4, earth model D-20time (kyr BP)

-20

-20

time (kyr BP)

point 4, earth model C
isostatic deformation (m)isostatic deformation (m)00time (kyr BP)

point 4, earth model A-20

10 20 30 40 50 60

isostatic deformation (m)-40isostatic deformation (m)-200point 1, earth model D

10 20 30 40 50 60

time (kyr BP)

point 6, earth model C

point 6, earth model D

0

0

-20
-40
-60
-80
-100
-120
0

10 20 30 40 50 60

isostatic deformation (m)0point 1, earth model C
isostatic deformation (m)isostatic deformation (m)point 1, earth model A

time (kyr BP)

-20
-40
-60
-80
-100
-120
0

10 20 30 40 50 60

time (kyr BP)

Key: Loading Models
Double80%

Standard
40 ka

Figure 3-40. Isostatic recovery sensitivity tests. Predictions of deformation between the present day and60 kyr in the future. Deformation at the present day may be regarded as the remaining deformation before
equilibrium is attained. Other details as for Figure 3-39.

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Table 3-6. Results of the isostatic recovery sensitivity tests. Each triplet of numbers, ’(a/b) c’,describe the time of maximum depression in kyr BP (a), the time of return to equilibrium in kyr
after present (b), and the time it takes to return to equilibrium in kyr (c = b–a).
Earth Loading
model model

Point 1

Point 2

Point 3

Point 4

Point 5

Point 6A

40 kyr

(–14.5/28) 42.5(–14.5/28) 42.5 (–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59C40 kyr

(–14.5/28) 42.5(–14.5/28) 42.5 (–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59D

40 kyr

(–14.5/36) 50.5

AStandard (–14.5/28) 42.5(–14.5/36) 50.5 (–14.5/36) 50.5 (–14.5/40) 54.5 (–15/44) 59(–17/56) 73(–14.5/28) 42.5 (–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59

C

Standard (–14.5/28) 42.5(–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59

D

Standard (–14.5/36) 50.5(–14.5/36) 50.5 (–14.5/36) 50.5 (–14.5/40) 54.5 (–15/44) 59

A

80%(–14.5/28) 42.5 (–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59(–14.5/28) 42.5(–17/52) 69

C80%(–14.5/28) 42.5(–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59

D

80%(–14.5/36) 50.5(–14.5/36) 50.5 (–14.5/36) 50.5 (–14.5/40) 54.5 (–15/44) 59

A

Double(–14.5/28) 42.5(–14.5/28) 42.5 (–14.5/28) 42.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59(–17/56) 73

C

Double(–14.5/28) 42.5(–14.5/32) 46.5 (–14.5/32) 46.5 (–14.5/32) 46.5 (–14.5/36) 50.5 (–15/44) 59

D

Double(–14.5/36) 50.5(–14.5/36) 50.5 (–14.5/36) 50.5 (–14.5/40) 54.5 (–15/44) 59(–17/56) 73

Following unloading, the time to reach equilibrium varies between ~40 kyr and ~75 kyr. This range,
and the observation that regions of maximum deformation recover more rapidly, may partly be attributed to the non-linear nature of solid Earth rebound. However, the distribution of relaxation times isprincipally an artefact of the method by which rebound is calculated: The rebound criterion states that
deflection must return to within 1% of the maximum deflection. Therefore regions of greatest deformation seem to recover more rapidly as they only need return to a larger absolute deflection. The positive
isostatic deformation at point 6 (Figure 3-39) indicates that this site was uplifted as part of a glacial
forebulge prior to being loaded. For all sites, the time for recovery was found to be similar for Earthmodels A and C, but the model with the stiffer lower mantle (Earth model D) took consistently longer
to reach maximum deformation and recover to equilibrium.
The ‘80%’ model was included in this study because it gives the closest fit to present-day uplift rates
using the 1D GIA model. However, despite attaining a smaller maximum amount of deformation, thelong-term rate of recovery for this loading model is not substantially different from that for models
with 100% ice thickness. Therefore, the rate of isostatic recovery is not strongly dependent upon themagnitude of the load, but rather the timing of loading.
Importance of far-field ice sheets

Variations in the far-field Laurentide and Antarctic ice sheets perturb the shape and height of the global
geoid signal, and generate an isostatic solid Earth response. Due to the distance of Fennoscandia fromthese ice sheets, the geoid signal will effectively be detected as a near-uniform rise or fall in the heightof the geoid/ocean surface, dependent upon the surface mass change associated with these far-field
ice sheets and distance from these surface masses. The solid surface signal due to far-field ice sheets
will only be detectable in Fennoscandia if this region is located upon the forebulge of the deformation, as was the case during the maximum extent of the Laurentide ice sheet during the LGM/Mitrovica et al. 1994/.
The magnitude of the solid Earth signal due to far-field ice sheet loading was calculated using astandard global ice loading model for a single glacial cycle /Tushingham and Peltier 1991/ with theice removed from Fennoscandia during the last glacial cycle. Estimates of horizontal and verticaldeformation at the proposed site of Forsmark are shown in Figure 3-41.
The positive sign of the predicted vertical deformation due to far-field ice loading at Forsmark(Figure 3-41a) indicates that the site was situated on the deformational forebulge of the Laurentide ice
sheet, as predicted by /Mitrovica et al. 1994/. However, the calculated signal may overpredict reality,
due to the neglect of any lateral structure in the Earth model, see section A GIA case study with a
Fennoscandian 3D Earth structure below. The negative sign of the horizontal deformation indicatesmotion in an easterly direction. The non-zero magnitude of predicted deformation at the present day
simply implies that the system has not yet returned to equilibrium following the removal of the
far-field ice load.
96

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25

horizontal deformation (m)vertical deformation (m)Isostatic deformation due to Far Field ice sheets
(a)

20
15
10
5
0
-100

-75

-50

-25

time (kyr BP)

0

4
2

(b)

Key: Earth models

0

A

-2

B

-4

C

-6
-8

D

-10

E

-12
-100

-75

-50

-25

0

time (kyr BP)

Figure 3-41. Solid surface deformation at Forsmark due to far-field ice loading. See Table 3-5 for a
description of the Earth models used. a) vertical deformation, b) horizontal deformation.

The vertical deformation peaked at 18kyr BP, coincident with the maximum extent of the Laurentide
ice sheet. The steady increase in the signal over time is a result of the gradual growth of the Laurentide
ice sheet throughout the last glacial cycle. The magnitude of this deformation reached ~25 m for
Earth models with a smaller lower mantle viscosity of 1·1022 Pa s (models A, C and E), and only
~20 m for Earth models with a stronger lower mantle (models B and D). The strong lower mantle
models demonstrate a slower, smaller response to loading, and a slower rate of recovery following
unloading.
This separation in the style of deformation according to lower mantle viscosity is also seen in thesmaller horizontal signal (Figure 3-41b). When a stronger lower mantle is assumed, the magnitude of
deformation is smaller during the majority of the glacial cycle, except during deglaciation, when therecovery of the solid Earth is delayed by ~10 kyr for these models. The reason for this is unclear, butshould provide a useful constraint upon lower mantle viscosities when testing predictions against data.
In general, horizontal deformation peaks later than vertical deformation; at ~14.5 kyr BP for models
with a weaker lower mantle, and ~4 kyr BP for models with a stronger lower mantle. A greater
magnitude of horizontal deformation is predicted for models with a thinner lithospheric thicknessof 96 km (as opposed to 120 km), but the absolute magnitude of the difference is small, and thisphenomenon is not apparent in the vertical signal.
The magnitude of the far-field signal is small in comparison with predictions of near-field isostatic
deformation (less than 10% of the total signal) during the Weichselian glaciation, but must not be
neglected.
A GIA case study with a Fennoscandian 3D Earth structure

Note that in this 3D Earth structure case study, adapted from /Whitehouse 2009/, the overall aim is
to investigate the sensitivity of GIA predictions to lateral Earth structure. To this end a comparison ismade between dedicated 1D and 3D GIA simulation where the Earth structure has been specificallyadjusted in order to give a relevant comparison between the two models. In turn, this means that the
Earth structures have not been selected to give a best fit of e.g. uplift rates and shore-level displacement against observed data. Also note that the 1D-3D comparison study described here uses anotherice sheet load history than the Weichselian glacial cycle used in the 1D simulations described inprevious parts of Section 3.3.4.
State-of-the-art GIA models that account for 3D Earth structure use a range of geophysical observables
to infer lateral variations in lithospheric thickness and mantle viscosity. Rayleigh wave dispersion
data have been used to indicate variations in the thickness of the seismic lithosphere in Fennoscandia/Calcagnile 1982/, while coherence studies have been used in the same region to study gradients offlexural rigidity and elastic thickness /Poudjom Djomani et al. 1999, Watts 2001, Pérez-Gussinyé
et al. 2004/. Fennoscandian lithospheric thickness variations, derived from elastic thickness
estimates, are shown in Figure 3-42.

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97

72N

32N

32E

42E
90

150

210

270

330

390

(km)

Figure 3-42. Lithospheric thickness derived from the elastic thickness model of /Watts 2001/ and /PérezGussinyé et al. 2004/. This is part of a global model where the thicknesses have been scaled to ensure a
global mean lithospheric thickness of 120 km. The red lines define plate boundaries. The dashed area isused to define a Fennoscandian regional mode studied in /Whitehouse 2009/ (see text).

A first-order observation from GIA studies that include 3D Earth structure is that lateral variations
in mantle viscosity affect the magnitude, but not the pattern, of GIA-related uplift rates, however
they affect both the magnitude and direction of tangential velocity rates /Kaufmann and Wu 1998,
Kaufmann et al. 2000, Kaufmann and Wu 2002, Latychev et al. 2005b, Whitehouse et al. 2006/.
A case study into the effect of including realistic Earth structure in Fennoscandian GIA models ispresented here.
The sea-level theory presented in /Mitrovica and Milne 2003/ is implemented using the 3D finite
volume model of /Latychev et al. 2005a/ to solve the sea-level equation on a spherical, rotating
Earth. Even though this study focuses upon Fennoscandia, global calculations are carried out toensure the correct redistribution of meltwater throughout the oceans.To give a global ice-loading
history for the whole of the last glacial cycle that is tuned to fit global ice volumes as determinedusing far-field sea-level data /Bassett et al. 2005/, the ICE-3G global ice model of /Tushingham andPeltier 1991/ is used for all ice sheets except in Fennoscandia where the ice model of /Lambeck et al.
1998/ is used, see Figure 3-43.
180°

24

0°

0°
12

60°

0°
30

0°
0

600

1200

1800

2400

3000

3600

Ice thickness (metres)

Figure 3-43. Ice thicknesses in the Northern Hemisphere at the LGM (20 kyrs BP). From the model of/Tushingham and Peltier 1991, Lambeck et al. 1998/.

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Calculations are initially carried out using a 1D Earth model with a lithospheric thickness of 120 km,
an upper mantle viscosity of 5·1020 Pa s, and a lower mantle viscosity of 5·1021 Pa s. The calculations
are then repeated using an Earth model that includes lateral variations in lithospheric thickness andmantle viscosity, and attempts to account for the presence of plate boundaries. The 3D Earth model
was derived using estimates of elastic thickness as a proxy for lithospheric thickness, and seismic
velocity anomalies as a proxy for mantle viscosity. Plate boundary zones are defined to be 200 kmwide bands of low (2·1020 Pa s) viscosity within the lithosphere. Portions of this model are shownin Figure 3-42.
A scaling method is used to ensure that the 1D and 3D Earth models have the same depth-averaged
mean structure; thus permitting a meaningful comparison of the results. For the lithosphere, thismeans that the elastic thicknesses of the 3D model were multiplied by a uniform constant to ensurethat the global mean lithospheric thickness is still 120 km. This results in the use of lithosphericthicknesses of more than 300 km for the Fennoscandian shield: these values are larger than independent
estimates e.g. /Calcagnile 1982/, but it should be noted that this study was carried out to explore thesensitivity of GIA calculations to variations in lateral Earth structure, and does not seek to accurately
replicate global 3D Earth structure.
For the mantle, the logarithm of the lateral viscosity perturbations at each depth (the calculations arecarried out upon a finite volume grid) are multiplied by a uniform constant to ensure the mean forthat depth is the same as in the 1D model. For further details of the construction of the Earth modelsee /Whitehouse et al. 2006/.
Figure 3-44 shows the predicted present-day uplift and horizontal deformation rates for Fennoscandia
generated using the 1D Earth model. Uplift is centred upon the region of maximum ice-loading, andhorizontal deformation exhibits a radial pattern of motion away from the centre of uplift.
Figure 3-45 shows the misfit in present-day vertical and horizontal deformation rates between the
1D and 3D models. Note that the maximum differences of ~3 mm/yr for uplift rates and ~1 mm/yr
for horizontal rates are greater than the current limit of observational uncertainty (e.g. from GPS
measurements), implying that any inferences of Earth structure obtained from the inversion ofgeodetic data will be biased if a 1D profile is assumed.
In Figure 3-45 the 1D model predictions are subtracted from the 3D model predictions, therefore
the blue region surrounded by the orange band in Figure 3-45a indicates that the 1D model predicts
higher maximum uplift rates and lower minimum subsidence rates (more negative) than the 3Dmodel. The differences are predominantly related to the magnitude of the predictions, and not thepattern, indicating that the introduction of lateral Earth structure has little effect upon the geometryof predicted uplift rates. However, in Figure 3-45b the change in the direction of the arrows indicatesthat both the magnitude and the pattern of horizontal deformation are perturbed with the introductionof realistic lateral Earth structure.
a)

b)

72N

72N

a)

b)

52N

52N
42E
2E

2E
–3

0

3

6

2 mm/y42E

9 12 (mm/y)

Radial Motion (mm/y)

Lateral Motion

Figure 3-44. Predicted present-day uplift rates (a) and tangential rates (b) generated using the best-fitting
1D Earth model for Fennoscandia and the ice model discussed in the text. Tangential rates are given relative
to a point on the north coast of the Gulf of Bothnia.

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72N

a)

b)

52N
2E

42E 2E

42E
1 mm/y–3.00–2.33 –1.65 –0.975 –0.3

0.375 1.05

Radial Motion (mm/y)

Lateral Motion

Figure 3-45. Misfit between (a) uplift rates and (b) tangential rates generated using a 3D Earth model anda 1D Earth model. 1D rates are subtracted from 3D rates.

The question of whether it will be possible to constrain 3D Earth structure via GIA modelling remains
open. The large number of degrees of freedom means that it is unlikely to be possible to carry out
a unique inversion of 3D viscosity structure using GIA observables. /Wu et al. 2005/ claim that 3D
Earth structure can be resolved using relative sea-level data from the centre of present-day rebound.
However, this is disputed in other studies e.g. /Spada et al. 2006, Whitehouse et al. 2006/, due to thesimilarity of the effect of lateral structure at different depths upon GIA observables. We investigate this
aspect of GIA modelling below.
Figure 3-46 again shows the difference (3D minus 1D) between predicted vertical (left) and horizontal(right) deformation rates. A different aspect of the 3D model is varied in each row. In the first row(Figure 4-46a) lithospheric thickness variations are permitted in the 3D calculations, but upperand lower mantle viscosities are held at the uniform values of the 1D model. In the second row(Figure 3-46b) a band of low viscosity is defined within the lithosphere along plate boundaries toreplicate the weak stress-coupling across these regions. Other parameters are fixed at the 1D values.
In the third row (Figure 3-46c) upper mantle viscosities are allowed to vary laterally whereas the lower
mantle viscosity and lithospheric thickness are held fixed, while in the fourth row (Figure 3-46d) lower
mantle viscosities are allowed to vary while the upper mantle viscosity and lithospheric thickness are
fixed. In the first, third and fourth cases, weak plate boundary zones are not included.
The similarity of Figures 3-46a and 3-46c, both in the magnitude of uplift differences and the directionof the horizontal misfit arrows, indicates that lateral Earth structure at different depths can produce
similar velocity perturbations in Fennoscandia. Therefore it is unlikely that observations of present-day
deformation rates will be able to uniquely infer the underlying Earth structure, or indeed identify thedepth of any lateral heterogeneity.
The pattern in Figure 3-46d does not replicate the distribution of LGM ice-loading as in the previous
cases, indicating a lesser dependence upon local ice-loading at lower mantle depths. Due to thetransmission of stress, the lower mantle in Fennoscandia will play a part in supporting ice-loading inNorth America as well as Fennoscandia, and the pattern of horizontal perturbations, directed towards
North America, reflects this.
The introduction of weak plate boundary zones induces a significant perturbation in horizontal deformation rates, but not vertical rates (Figure 3-46b). The reason for this is that the band of low viscosity
decouples the transfer of stress across the plate boundary; the direction of the misfit arrows again reflects
the non-negligible effect of North American ice sheet loading upon surface deformation in Europe.
Uplift rates have a strong local dependence upon loading; hence there is no perturbation to these rates
when plate boundaries are introduced.
The effect of introducing lateral Earth structure at different depths has an approximately linear
cumulative effect upon present-day deformation rates. This may be deduced by observing that within
the region of uplift the sum of the perturbations in Figure 3-46 approximately equals the cumulative
effect (Figure 3-45a).

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72N

a)
52N
72N

b)
52N
72N

c)
52N
72N

d)
52N
2E

42E

–1.5 –1.2 –0.9 –0.6 –0.3 0

0.3 0.6 0.9 1.2 1.5

Radial Motion (mm/y)

2E

42E
1 mm/yLateral Motion

Figure 3-46. (a) Misfit due to lateral variations in lithospheric thickness. (b) Misfit due to the inclusion of
weak plate boundary zones. (c) Misfit due to lateral variations in upper mantle viscosity. (d) Misfit due tolateral variations in lower mantle viscosity. Left column: difference in uplift rates. Right column: differencein horizontal rates. All plots show 3D predictions minus 1D predictions.

Predictions of relative sea-level history for the last 10 kyrs are also compared for the 1D and 3Dmodels. Figure 3-47a shows the perturbation to relative sea-level predictions at 10 kyrs BP due tothe inclusion of lateral Earth structure (3D prediction minus 1D prediction). There is again a strong
correlation between the pattern of ice loading and the region of maximum perturbations, and there isalso a large positive misfit to the east of Fennoscandia in the region of maximum lithospheric thickness(Figure 3-42). Once again the magnitude of perturbations due to the introduction of lateral Earthstructure is greater than the observational uncertainty of relative sea-level data, which can be up to
10m. Therefore, the results show that inverting relative sea-level data to yield inferences of Earthparameters or ice loading history will lead to biased results if lateral structure is neglected.
Along the Norwegian west coast there is very little difference in predictions of relative sea-level by
the two models at 10 kyrs BP. In the top left plot of Figure 3-48 the relative sea-level history for acoastal site with good relative sea-level data (Bjugn: shown by the left-hand star in the bottom plotof Figure 3-48) is plotted for the whole of the Holocene. The black curve is the prediction fromthe 1D model and the red curve is the prediction from the 3D model. The similarity of the curves
is characteristic for sites along the length of this coastline, indicating that introducing lateral Earthstructure into the GIA model makes little difference to Holocene relative sea-level predictions forthis region. Relative sea-level data from the Norwegian west coast may therefore form a data set thatis insensitive to the 3D Earth structure of this region, and hence may be used to tune GIA model inputs
such as 1D Earth structure and ice load history. In this case, the poor fit to the data may be rectified
by altering either the ice history or the depth-averaged viscosity profile of the Earth model.
The geographical region over which such data would be able to constrain input parameters is limited;
it would be useful to also use relative sea-level data from within the Gulf of Bothnia to tune the GIA
model. The predicted relative sea-level history for Ångermanälven is shown in the top right plot of
Figure 3-48. Ångermanälven is the site of an extensive Holocene relative sea-level data set. It is situatedon the east coast of Sweden (right-hand star in the bottom plot of Figure 3-48), in the centre of theregion of maximum misfit for both uplift rates and 10 kyrs BP relative sea-level. The large difference
between the two curves during the last 10 kyrs indicates that using such data to tune GIA model inputs
will lead to biased inferences of Earth structure and ice history.

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a)

75

b)

75

75

70

70

70

70

65

65

65

65

60

60

60

60

55

55

55

55

50

50

50

50

45

45

45

45

40

40

40

40

35

35

35

0

20

75

35

40

20

0

–70 –60 –50 –40 –30 –20 –10 040

10 20 30 40 50 60 70

rsl (m)Figure 3-47. Difference in relative sea-level predictions at 10 kyrs BP. (a) 3D model minus 1D model.
(b) 3D model minus ‘regional’ 1D model (see text).

Bjugn

120
100

rsl (m)rsl (m)

80
60
40
20
0

–11 –10 –9 –8 –7 –6 –5 –4 –3 –2 –1

0

260
240
220
200
180
160
140
120
100
80
60
40
20
0

Ångermanälven

–11 –10 –9 –8 –7 –6 –5 –4 –3 –2 –1

time (ka)0time (ka)

72N
Bjugn

Ångermanälven
52N
2E

42E

Radial Motion (mm/y)–3.00 –2.33–1,65 –0.975 –0.3

0.375

1.05

Figure 3-48. Top left: Predicted relative sea-level history at Bjugn for the last 10 kyrs. Black line: 1D model.
Red line: 3D model. Top right: Predicted relative sea-level history at Ångermanälven for the last 10 kyrs.
Line colours are as for Bjugn. Relative sea-level data are shown as triangles; see /Lambeck et al. 1998/ forreferences. Bottom plot: Location map for Bjugn and Ångermanälven overlaid upon the misfit in uplift rates(as for Figure 3-45a).

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The difference in relative sea-level predictions for the 1D and 3D models is also plotted for Oskarshamn
and Forsmark in Figure 3-49. To exemplify the magnitude of the error introduced by using a 1D Earthmodel instead of a 3D Earth model to predict relative sea-level, as done in the GIA studies, the difference
at 10 kyrs BP is ~60 m at Forsmark (Figure 3-49).
The difference between 1D and 3D GIA simulations presented in this case study implies that ignoringlateral Earth structure and using a 1D Earth model in Fennoscandia leads to an over prediction ofpresent-day uplift rates and the total amount relative sea-level change during deglaciation. This, inturn, conclusively shows that a significant part of the misfit between the 1D GIA model results andGPS observations arises from the 1D Earth structure assumption. In addition, the misfit could also
include a component of having too large ice thickness in the Weichselian ice sheet reconstruction. Inline with this, one general interpretation of the results from the case study is that using a 1D model
will lead to inferred LGM ice thicknesses that are too small. A second general interpretation is thatthe use of a 1D model will lead to overestimates of mantle viscosity or lithospheric thickness in an
attempt to fit the observational GPS and relative sea-level data.

240
220
200
180
160
140
120
100
80
60
40
20
0

140

Forsmark

Oskarshamn

120
100

rsl (m)rsl (m)

It is computationally demanding to solve the sea-level equation using a full 3D model; therefore
there is ongoing work to determine to what extent we can continue to improve our understandingof GIA using flat Earth models or a suite of spherically-symmetric 1D models. /Paulson et al. 2005/
suggest that GIA observables within formerly-loaded regions are only sensitive to the local viscosity
structure, and a global model that adopts the relevant local viscosity structure should be sufficient tomodel GIA in that region. The situation becomes more complex away from the centre of the former
ice sheets because GIA in these regions seems to be dependent upon both the local viscosity structure
and that beneath the ice sheets, which may be different, thus a model that allows for lateral variations
in viscosity is required.

80
60
40
20
0

–11 –10 –9 –8 –7 –6 –5 –4 –3 –2 –1 0

time (ka)5°–11 –10 –9 –8 –7 –6 –5 –4 –3 –2 –1 0

10°

15°

Forsmark

60°

time (ka)

20°

60°

Oskarshamn

55°

5°

10°

15°

20°

55°

Figure 3-49. Predicted relative sea-level curves for Forsmark and Oskarshamn calculated using 1D (blackline) and 3D (red line) Earth models. Note that both the 1D and 3D GIA simulations used for this plot
use another ice load history than the 1D GIA simulations in the previous parts of Section 3.3.4. The
results should only be used to exemplify the error introduced by using 1D instead of 3D GIA modelling.
At Forsmark, the difference at 10 kyrs BP is ~60 m.

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3.3.5 Time perspective
Shoreline migration is an ongoing process. It is strongly coupled to the global distribution of water
between the oceans and land-based ice sheets, and the isostatic response of the Earth to loading
during a glacial cycle. This isostatic deformation is a slow process in relation to the rate at which ice
sheets grow and decay during such a cycle.
As ice sheets grow, water is removed from the oceans and global eustatic sea-level falls. During thebuild-up to the last glaciation, eustatic sea-level fell at a mean rate of ~1.3 mm/yr between 120 kyr
BP and 20 kyr BP /Fleming et al. 1998, Yokoyama et al. 2000, Clark and Mix 2002, Milne et al. 2002,
Peltier 2002, Mitrovica 2003/, although there were many departures from this rate during the build-up
period, due to higher frequency climatic oscillations /Imbrie et al. 1984, Lambeck et al. 2002,
Waelbroeck et al. 2002, Siddall et al. 2003/. During deglaciation, the initial rate of eustatic sea-level
rise between ~21 kyr BP and 17 kyr BP was only ~6 mm/yr, followed by an average rate of ~10 mm/yr
for the next 10 kyr /Fleming et al. 1998/. Perturbations from this rate occurred during meltwater pulses
at ~14.5 kyr BP and ~11 kyr BP. These are attributed to rapid melting events, during which rates of
eustatic sea-level rise reached ~15 mm/yr /Fairbanks 1989, Bard et al. 1990, 1996, Clark et al. 1996,
Fleming et al. 1998, Lambeck and Chappell 2001, Lambeck et al. 2002, Peltier 2005, Bassett et al.
2005/. Rates of eustatic sea-level change fell to negligible values following the end of deglaciation
around 5 kyr BP. However, meltwater is not distributed uniformly throughout the oceans due to thealteration to the shape of the geoid following the redistribution of surface masses. Sea-level can actually fall at sites within ~20° of a melting ice sheet /Farrell and Clark 1976/, therefore the rate of eustatic
sea-level change is a poor indicator of local rates of sea-level change, especially for a local melt source.
The local rate of relative sea-level change incorporates both local changes in sea-level and the isostatic
deformation of the solid Earth. The balance between rates of isostatic deformation and local sea-levelchange determine whether a site has advancing or retreating shorelines.
During the final build-up of the Fennoscandian ice sheet from ~32 kyr BP, isostatic rates of solid
Earth subsidence due to surface loading by ice are estimated to have reached values of 40 mm/yr (seeSection 3.3.4). During future glacial cycles, modelling predictions imply that maximum rates will beattained immediately prior to the time of greatest ice thickness. In general, the time of maximum ice
thickness is predicted to precede the time of maximum solid Earth deformation, with the latter being
delayed by up to ~3 kyr. During the period between maximum ice thickness and maximum deformation, the rate of solid Earth subsidence will decrease to zero.
There is usually a delay between the time of maximum deformation and the time at which a location
becomes ice-free. Again, modelling predictions imply that as an ice sheet thins, rates of isostatic
rebound will be low; < 20 mm/yr for this reconstruction of the Weichselian. Maximum rebound rates
occur immediately following the final removal of an ice sheet, with uplift rates reaching ~75 mm/yr
in this model at sites which had the greatest ice cover. The rate of rebound then decays exponentiallywith time; maximum rates at the centre of present-day uplift in Fennoscandia are ~10 mm/yr, andthese are expected to decay to negligible values within the next 30 kyr (see Section 3.3.4).
During a glacial cycle, immediately following the removal of ice, local uplift rates are far greater than
the rate of local sea-level changes, consequently a situation of regression dominates, and shorelines
migrate oceanwards. The rate of shoreline migration depends upon the local gradient of topography.
A typical Swedish Baltic Sea shoreline with a gradient of 2 m of elevation per 1km would experience
35 m/yr of oceanward shoreline migration for an uplift rate of 70 mm/yr, or 5 m/yr of oceanward
shoreline migration for an uplift rate of 10 mm/yr (the present maximum in the Baltic region),assuming negligible changes in local sea-level.
A combination of decreasing rates of isostatic rebound and sea-level rise due to far-field melting
slows the rate of regression with time, with a switch to transgression taking place if rates of local
sea-level rise exceed the rate of local isostatic rebound. This situation is likely to arise during a
global meltwater pulse at sites with low rebound rates, but will not be maintained because melting
takes place on a shorter time scale than isostatic rebound. Rebound will once again become thedominant factor governing shoreline migration as sea-level changes become negligible. During thecurrent interglacial period decaying uplift rates persist. These, along with processes related to ocean
syphoning (see Section 3.3.1), act to maintain a situation of gradual oceanward shoreline migration
throughout the majority of the Baltic region for the first few 10 kyr of an interglacial so long as
the higher interglacial temperatures do not give rise to a longer-term melting event, such as thedestabilization, and subsequent melting, of the Greenland Ice Sheet.
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3.3.6 Handling in the safety assessment SR-Site
Isostatic evolution and shoreline displacement have been investigated by means of numericalmodelling. In this context, the main question for the safety assessment is whether the Forsmark site
is submerged or not. The salinity in the lake/sea covering the site is also of interest. A 1D GIA model
has been used to simulate shoreline migration. The evolution of the shoreline has been calculated fora glacial-interglacial cycle using a loading function based on the Weichselian ice sheet simulation
described in Section 3.1.4. The GIA model used for generating the Weichselian glacial cycle is thatdeveloped by Milne /Mitrovica and Milne 2003/. The global ice-loading function used in the study
is modified from the ICE3G deglaciation history /Tushingham and Peltier 1991/, and has beencalibrated using far-field relative sea-level data. A eustatic curve has been used to tune the massof ice contained within far-field ice sheets. The Earth model is based on Maxwell rheology with a
1D radial three-layer structure. The lithosphere is represented by a 96 km thick layer with a very
high viscosity and thus behaves as an elastic medium over GIA time scales, the density and elastic
structure are from the preliminary reference Earth model (PREM) which has been determined to ahigh degree of accuracy by seismic methods /Dziewonski and Anderson 1981/.
At the start of a GIA model run, the Earth is assumed to be in isostatic equilibrium. In reality, theEarth is unlikely to reach such a state if glaciations occur with similar periodicity as in the past. To
correct for this, the GIA modelling has been initiated by a glacial loading history yielding shorelines
comparable to those reported by /Funder et al. 2002/ at the peak of marine inundation in the Early
Eemian about 130 kyrs before present.
Complementary 3D GIA simulations have been made in order to quantify the error introduced byusing a laterally homogeneous Earth model. The results show that the 1D modelling simulations
probably over-predict the isostatic response to the ice sheet load as well as present day uplift rates.
For the analysis of the biosphere and hydrological evolution the shoreline evolutions during the firstthousands of years of the initial period with temperate climate domain is extrapolated from shoreline
data /Påsse 2001/, see Section 4.5.2. For further, detailed information on the GIA modelling, see thepreceding parts of Section 3.3.

3.3.7

Handling of uncertainties in SR-Site

Uncertainties in mechanistic understanding

The processes involved in GIA, and their effect upon shoreline migration, are very well understood.
There are no major uncertainties in our understanding of the mechanistic processes that cause
shoreline migration.
Model simplification uncertainty•

Discussions of the uncertainty in the calculated shoreline are based on the sensitivity analysis andcase study presented above.
The assumption of a 1D Earth model is a simplification of the situation in Fennoscandia. A com­
parison between 1D and 3D GIA simulations /Whitehouse 2009/ shows that the assumption ofa laterally homogeneous (1D) Earth structure over Fennoscandia probably results in an overprediction of the isostatic response to ice sheet load and present-day uplift rates. As an example,
the 1D-3D GIA model comparison shows that the 1D simulation potentially overestimates therelative sea-level change over the last 10 kyrs by up to 60 m at Forsmark.

•The time discretisation of the basic loading model is fairly coarse (every 7 kyr) between 116 kyr BP
and 21 kyr BP, reflecting the lack of constraints upon ice history and sea-level prior to the LGM. Iflong time steps are used the loading function may be under- or over-estimated, leading to inaccuracies in relative sea-level, and hence shoreline migration, predictions for this period. A higher time
resolution has been investigated, and discrepancies in relative sea-level predictions may be up to
50 m between a model that uses a time step of 7 kyr and one that uses a time step of 1 kyr, during
periods of rapid change in ice sheet geometry. However, using a coarse time scale during the earlystages of a model run has a negligible effect upon predictions of relative sea-level change during
the latter stages of a model run. A high resolution time scale has been used for ice loading since
the LGM, i.e. for the period when ice history is known reasonably accurately.

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•At the beginning of a model run, the Earth is assumed to be in a state of isostatic equilibrium,with no deformation remaining from a previous glacial cycle or any other process that upsetsthe isostatic balance of the system. In reality, as the Earth enters each glacial cycle it will be
pre-stressed, with deformation ongoing as a result of the previous loading event. GIA-induced
deformation decays over a few tens of thousands of years, and after this period there will be anegligible difference between predictions from a model that started from equilibrium and onethat was pre-stressed. However, in order to reproduce the relative sea-level evolution for the
Weichselian as closely as possible the model was run twice in succession, with the second run
taken to represent the Weichselian development. Nevertheless, a degree of inaccuracy remains at
early times in the model since deformation will depend upon the details of the Saalian deglaciation,
about which very little is known.

Input data and data uncertainty

As demonstrated by the similarity of the shoreline predictions for a range of rheological parametersin Section 3.3.4, uncertainties in the lithospheric thickness and mantle viscosity parameters are currently significant. The values of such parameters will be more tightly constrained following detailed
analysis of the results of the GIA modelling in comparison with relative sea-level and GPS data.

3.3.8 Adequacy of references
The SKB report produced for the handling of isostatic adjustment and shore-level changes /Whitehouse
2009/ has undergone the SR-Site QA system handling, including a documented factual review procedure.
Also the SR-Can Climate report /SKB 2006a/, from which some of the studies are used, has undergone
QA system handling including a factual review process. Other references used for the handling of ice
sheet dynamical processes are peer-reviewed papers from the scientific literature.

3.4

Permafrost

3.4.1 Overview/general description
Permafrost is defined as ground which remains at or below 0°C isotherm for at least two consecutive
years /French 2007/. This definition is based exclusively on temperature, and disregards the texture,degree of compactation, water content, and lithologic character of the material. As a result, the term
permafrost does not always equate to perennially frozen ground, since depending on pressure andcomposition of groundwater and on adsorptive and capillary properties of ground matter, water inground may freeze at temperatures below 0°C. Therefore, permafrost encompasses the perenniallyfrozen ground and a surrounding so-called cryopeg, i.e. a ground layer in which water remains
unfrozen at sub-zero temperatures.
Permafrost and perennially frozen ground originate from the ground surface depending on a complex
heat exchange process across the atmosphere/ground boundary layers and on an almost time-invariant
geothermal heat flow from the Earth’s interior. The heat exchange between the atmosphere andthe Earth’s surface is governed by solar radiation, longwave terrestrial and atmospheric radiations,sensible and terrestrial heat fluxes, and evaporation and condensation /Lockwood 1979, Washburn
1979, Lunardini 1981, Williams and Smith 1989, Smith and Riseborough 1996, Yershov 1998/.
Freezing of water within permafrost is based on the thermodynamics of interdependent thermal,
hydraulic, mechanical, and chemical processes in the ground. In addition, freezing of water influences the thermal, hydraulic, mechanical, and chemical behaviour of the ground. Thermal properties
change from those of unfrozen ground to those of frozen ground, affecting the heat transfer process.
Ice formation in pores of the ground confines groundwater flow through the almost impermeable
frozen ground, therefore altering the overall groundwater circulation. Ice formation can also cause
deformation of the ground and changes in the mechanical stress state. Frost weathering and degradation of the ground surface and patterned ground are additional consequences of cyclic freezing andthawing processes. Moreover, exclusion of salts in the freezing of saline groundwater can lead toincreased salinity concentrations in the unfrozen cryopeg.

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Issues associated with permafrost development in the geosphere are comprehensively explainedin /Washburn 1979, Williams and Smith 1989, Yershov 1998, French 2007/ and reviewed in e.g.
/Gascoyne 2000, Ahonen 2001, Vidstrand 2003/. Experimental investigations of thermal-, hydrochemical- and mechanical impacts of freezing on bedrock properties have been reported by e.g. /Mackay
1997, Ruskeeniemi et al. 2002, Ruskeeniemi et al. 2004/, whereas model studies dealing with permafrost development in ground under climate change can be found in e.g. /Lunardini 1995, Kukkonen andŠafanda 2001/. More recently, some model studies on thermo-hydro-mechanical impacts of freezing
processes on bedrock properties with implications for interactions between glaciers and permafrost
in a time frame of a glaciation cycle (~100,000 years) have been conducted, e.g. /Bauder et al. 2003,
Hartikainen 2004, SKB 2006a, Person et al. 2007, Lemieux et al. 2008a, b, c/. The effects of freezingof the geosphere on groundwater flow have been studied by e.g. /Vidstrand et al. 2010/.

3.4.2

Controlling conditions and factors

Climate and surface conditions

The principal factors controlling the heat exchange between the atmosphere and the Earth’s surface
are climate, topography, vegetation and snow covers, soil characteristics and water bodies. These
factors are mutually dependent and can vary considerably in time and location. The main climaticparameters are insolation, air temperature, wind, and precipitation. Insolation is a driving force
governing the heat exchange between the atmosphere and the Earth’s surface and affecting the otherclimatic parameters. Air temperature, which is commonly applied to map permafrost distribution,
controls the longwave atmospheric radiation, the turbulent heat exchange, and evaporation andcondensation. Wind, in turn, mainly influences the sensible heat exchange, but also latent heat production and loss. Precipitation together with evaporation and condensation determine groundwater
recharge, affecting the groundwater content and flow and hence the terrestrial heat flux.
Topography has a significant impact on climate conditions. Generally, air temperature decreases
as altitude increases being affected by radiation, convection and condensation. The average airtemperature decrease is approximately 0.65°C for every 100 metres increase in height /Danielson
et al. 2003/. In addition, inversions are common in hilly terrain causing low-slope low-lying areas toexperience significantly lower temperatures than higher lying and steeper-slope areas. Furthermore,
the slope angle and azimuth affect the flux of shortwave radiation, and where topographical differences are large a more patchy distribution of permafrost is expected.
Vegetation and snow covers are sensitive to climatic conditions and topography. The characteristic
parameters of the surface cover are surface albedo, emissivity, and roughness controlling the incoming
shortwave radiation, the longwave terrestrial radiation and the turbulent heat exchange, respectively,
as well as the thermal properties and the thickness of the surface cover affecting terrestrial heattransfer. In general, the vegetation and snow cover moderate the ground temperature and thus theaggradation of permafrost. Vegetation is an insulating cover limiting cooling in winter and warmingin summer, hence reducing the annual fluctuation of ground temperature. Vegetation is also important
for the creation of snow cover, which protects the ground from heat loss in winter. However, the highalbedo of snow can lead to a snow surface temperature almost 2°C lower than the mean winter air
temperature /Yershov 1998/. As a rule, for a majority of the surface covers, permafrost can buildfrom the ground surface if the annual mean air temperature is lower than a value ranging between–9 and –1°C /Washburn 1979, Williams and Smith 1989, Yershov 1998, French 2007/. An exception
occurs with peat layers, which can insulate the ground from warming in summer more effectively
than from cooling in winter, with the resulting effect that permafrost can exist where the meanannual air temperature is above 0°C /Williams and Smith 1989/.
Properties and thickness of the soil cover affect the terrestrial heat flow. Of importance are the porosity
and water content of the soil, influencing the annual fluctuation of ground temperature and the thicknessof the active layer, i.e. the seasonally thawing ground layer. The soil cover also acts as an insulating
cover, since the thermal conductivity of the soil cover is lower than that of the underlying bedrock.
Water bodies, i.e. sea, lakes, and watercourses, influence permafrost creation and distribution considerably
since they have high specific heat content. A talik, i.e. an unfrozen layer, can exist beneath water bodies
that do not freeze to their bottom in winter. Depending on the characteristics of the climatic zone, the critical depth of a water body to remain unfrozen in winter is approximately 0.2 to 1.6 metres /Yershov 1998/.

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Sea water extensively reduces the development of costal permafrost. On the other hand, when the
shore-level of a highly saline sea is rising, submerged permafrost and perennially frozen deposits cansurvive for a long time beneath a cold seabed /Washburn 1979, Yershov 1998/.
The hydrological conditions on the ground surface affect the freezing of groundwater. Especially
under glaciated conditions, a warm-based overlying ice sheet increases the subglacial groundwaterpressure, in which case the freezing point may decrease to such a degree that the subglacial ground is
kept unfrozen. A similar, but minor, effect may occur when cold ground is submerged and submarine
freezing is reduced by the pressure of the sea water. Furthermore, groundwater flow, whether carrying
fresh glacial meltwater or saline seawater, can influence the freezing process by altering the chemical
groundwater composition.
Table 3-7 summarises how geosphere variables are influenced by permafrost development.
Subsurface and repository

The ground temperature that defines the presence of permafrost and primarily governs the freezingof water in ground is principally controlled by the ground thermal energy balance in terms of heattransfer, geothermal heat production, the specific heat content and the amount heat generated byphase change processes of water. Heat transfer within the ground can occur through conduction,
convection and radiation /Sundberg 1988/. In general, only conduction is regarded as important topermafrost evolution, since radiation is of importance only in unsaturated high-porosity ground athigh temperatures, and convection is of importance only when groundwater and gas fluxes are large.
Heat conduction depends on the ground temperature gradient, ambient temperature conditions, andthe thermal properties of the ground matter. Thermal conductivity, describing the ability of material
to transport thermal energy, and heat capacity characterising the capability of material to store heat,
Table 3-7. Influence of permafrost development on geosphere variables.
Geosphere variable

Climate issue variable

Summary of influence

Temperature

Permafrost depth

Permafrost is by definition a thermal condition
hence having no influence on temperature.

Groundwater flow

Frozen/unfrozen fraction of groundwater

Freezing occurs at temperatures below thefreezing point /SKB 2010c/. Groundwatertransformed from liquid to solid phase canbe regarded as immobile. Filling fracturesand pores, ice also affects the groundwaterflow through permeability, which decreases
with a reduction in the unfrozen fraction ofgroundwater.

Groundwater pressure

Frozen/unfrozen fraction of groundwater

The volume increase in the phase changeof water from liquid to solid state causes anincrease in the pressure of water which remains
liquid. Moreover, freezing of water in porous
ground matter can lead to cryosuction anddepression of groundwater pressure in theunfrozen ground.

Groundwater composition

Frozen/unfrozen fraction of groundwater

When saline water is transformed from liquid
to solid phase, the solutes are typically notincorporated in the crystal lattice of ice buttransferred in the liquid phase. Therefore thesalinity in the unfrozen fraction increases.

Rock stresses

Frozen fraction of groundwater

Deformations due to ice formation will leadto changes in rock stresses. The effect of thechanges is strongest nearest to the surface
where in situ stresses are lowest.

Fracture geometry

Frozen fraction of groundwater

Freezing can have some influence on thefracture geometry at shallow depths due tofrost cracking. Close to the ground surface
frost wedges can form to the depth of somemetre. Otherwise, there can be widening of
fractures due to freezing but the effect may bereversible as thawing takes place.

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depend on a number of variables such as mineralogy, porosity and groundwater content. Having typically three to four times higher conductivity than other common rock forming minerals, quartz is themost important mineral for determining thermal characteristics in native rock materials.
When the porosity is less than 1%, freezing of water has a minor effect on heat transfer in watersaturated ground. The degree of saturation has relevance to the thermal properties of ground due to thevery low thermal conductivity and heat capacity of air, e.g. the thermal conductivity of 1%-porosity
granite can decrease by over 10% with decreasing saturation /Clauser and Huenges 1995/. In rock,heat capacity is not very dependent on ambient temperature and pressure conditions, whereas thermal
conductivity is a rather variable function of both temperature and pressure. The thermal conductivityof granite decreases with increasing temperature by approximately 5 to 20% per 100°C and increases
with increasing pressure by about 1 to 2.5% per 100 MPa /Seipold 1995/. The pressure dependence ofthermal conductivity is increased when rocks are unsaturated /Sundberg 1988, Clauser and Huenges
1995/. Moreover, /Allen et al. 1988/ reported a strong correlation between lithology and permafrost
depth, which could be directly explained by differences in thermal conductivity.
In addition to the thermal properties above, the hydrogeochemical and mechanical properties of
the ground important for freezing of groundwater are permeability, porosity, adsorptive capacityof ground matter, chemical composition of groundwater and deformation properties of the ground.
Also, in the context of assessing the long-term safety of repositories for spent nuclear fuel, the spent
nuclear fuel will generate heat with an exponentially decreasing rate for a considerable amount oftime. This heat will warm up the surrounding bedrock /SKB 2010c/ and may thus act to reduce thethickness of permafrost above and near the repository (Section 3.4.4).

3.4.3 Natural analogues/observations in nature
At present, approximately 25% of the total continental land area of the Earth is occupied by permafrost.
Permafrost distribution may be characterised into: i) continuous (more than 90% spatial coverage),ii) discontinuous (between 90 and 50% coverage), iii) sporadic (less than 50% coverage), and iv)
subglacial forms. About one fifth of this permafrost is estimated to be subglacial in Antarctica andGreenland. Permafrost is abundant in Alaska, the northern parts of Canada and Russia, and in partsof China /French 2007/. Along the coast of southern and south-western Greenland both continuousand discontinuous permafrost can be found /Mai and Thomsen 1993/. The Northern Hemispheredistribution of permafrost is seen in Figure 3-50.
The deepest known permafrost occurs in the central part of Siberia in Russia, where thicknesses of up
to 1,500 m have been reported /Fotiev 1997/. The extensive region of continuous permafrost in central
Siberia corresponds to areas that are believed not to have been covered by Quaternary ice sheets andthat experienced cold subaerial climate conditions during the last glacial cycle and onwards. In coastal
areas, submarine permafrost may also exist, such as in northern Siberia where the permafrost formed
during ice free periods of the last glacial cycle in regions subsequently covered by the Arctic Ocean.
Furthermore, permafrost is frequently observed in mountainous terrain. For example, in the area ofTarfala in the Kebnekaise massif in northern Sweden, discontinuous permafrost has been reported tobe 100 to 350 m thick at an altitude above 1,500 m a.s.l. /King 1984, Isaksen et al. 2001/.
SKB, together with co-funding parties from Finland (Geological Survey of Finland and Posiva), GreatBritain (UK Nirex Ltd) and Canada (Ontario Power Generation and University of Waterloo), has carried out a research project on permafrost at the Lupin gold mine in northern Canada /Ruskeeniemi et al.
2002, 2004/. In this area, the depth of the permafrost extends to ~500−600 m. The main objective of the
project was to provide data describing the subsurface conditions in a permafrost area with crystalline
bedrock. Permafrost depth, temperatures, groundwater composition and hydraulic properties have beenmeasured. The mean annual air temperature at this site is −11°C and the annual mean precipitation is
low; ~270 mm /Ruskeeniemi et al. 2002/. Based on a seismic refraction survey, the depth of the active
layer has been interpreted to be ~1.5 m, varying between 1.2 and 1.8 m. The present permafrost depth
is believed to have been developed over the last 5,000 years. In the Lupin area, all shallow lakes freeze
down to the bottom during winter. Lakes deeper than 2–3 m are expected to have unfrozen bottoms
all year. The latter lakes may have the potential to support closed taliks. The large Lake Contwoyto,
located about 1,300 m from the mine at surface, provides the most significant talik structure that may
extend through the deep permafrost. However, direct observations beneath this lake are lacking.
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Figure 3-50. Permafrost distribution in the Northern Hemisphere. From /UNEP/GRID-Arendal 2005/.

In order to investigate the distribution of permafrost and locate salinity differences in the deep ground­waters, a number of electromagnetic soundings were conducted in the Lupin mine area. Furthermore, drillings were made from the deeper parts of the mine through the base of the permafrost, in order to sample
groundwaters and to study the distribution of open fractures and hydraulic conditions. The main resultof the electromagnetic sounding surveys was the identification of anomalies forming a subhorizontal
layer at depths between 400 and 700 m. /Paananen and Ruskeeniemi 2003/ made the interpretation that
this conducting layer represents saline or brackish water at the base of the permafrost. According to thetemperature measurements made in the mine, the base of the permafrost occurs at a depth of 540 m. The
drillings revealed the existence of a ~100 m thick unsaturated or dewatered zone below the permafrost.
An alternative interpretation of the subhorizontal conductor at ~650 m depth could be that it represents thegroundwater table /Ruskeeniemi et al. 2004/. The dry zone below the permafrost could either be a natural
result of very limited recharge through the permafrost or an effect of mine drainage. No pressurised
water or gas flow was observed in the boreholes. The available data from water sampling do not provide
any evidence of highly saline water below the permafrost /Ruskeeniemi et al. 2004/.

3.4.4 Model studies
Two climate cases have been studied by permafrost modelling; i) the Reconstructed last glacial cycle
and ii) the severe permafrost case, see Figure 1-3. In addition, a broad range of sensitivity tests havebeen made, covering the range of uncertainties associated with various parameters of importance forpermafrost development. The simulations of the Reconstructed last glacial cycle are described in thissection, whereas the simulations of the severe permafrost case are described in Section 5.5.
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The simulations of permafrost for the reconstruction of the last glacial cycle have been done with two
types of models, a 1D permafrost model and a 2D permafrost model improved from the 1D model. The
1D modelling was performed for the SR-Can safety assessment /SKB 2006a, Section 3.4/, but someof the results are used also for the SR-Site safety assessment. The 1D simulations were conducted
specifically for the repository location. The 2D modelling /Hartikainen et al. 2010/ was conducted for
the SR-Site safety assessment, and covers a 15 km long and 10 km deep vertical cross-section crossing
most of the Forsmark site. Both simulations use site-specific input data from the Forsmark site
investigation programme. Details on the input data used for the 1D and 2D permafrost simulations
are found in /SKB 2006a, Section 3.4/ and /Hartikainen et al. 2010/ respectively.
The purpose of the modelling studies is to analyse the main factors of importance for the developmentof permafrost and perennially frozen ground at Forsmark. The permafrost models include mathematicalexpressions for freezing and thawing of saline groundwater-saturated bedrock. The bedrock is con­sidered
as an elastic porous medium and the groundwater as an ideal solution of water and ionic solvents.
The models are based on the principles of continuum mechanics, macroscopic thermodynamics andthe theory of mixtures being capable of describing heat transfer, freezing of saline water, groundwaterflow and deformations of bedrock. In the 2D version of the model, the freeze-out and transport ofsolutes is included /Hartikainen et al. 2010/. The models are described further in /Hartikainen 2004,
SKB 2006a, Hartikainen et al. 2010/.
To capture the most important factors and parameters affecting the development of permafrost,
sensitivity analyses have been performed considering the following issues:•••

Surface conditions.
Subsurface conditions.
Presence of the repository.

Surface temperatures, together with the influence of surface covers such as snow, vegetation and water
bodies, have been included as factors of importance in the surface conditions. The investigatedbedrock conditions are thermal properties of the bedrock and the geothermal heat flow. The sensitivity
simulations carried out with the 2D permafrost model /Hartikainen et al. 2010/ are presented in thepresent report. Similar sensitivity experiments were carried out also with the 1D model /SKB 2006a,
Section 3.4/. However, since the sensitivity experiments conducted with the 2D model /Hartikainen
et al. 2010/ cover a broader range of cases than the sensitivity cases studied with the 1D model /SKB
2006a/, only some selected sensitivity tests from the 1D experiments are included in the present report.
The heat generated by the spent fuel has been included in most simulations, and is also analysed in a
dedicated 2D sensitivity simulation. In this context it should be noted that the repository depth waschanged between SR-Can (for which the 1D simulations were made) where it was 400 m and SR-Site(for which the 2D simulations were made) where it is 450 m. Despite of this the resulting permafrost
and freezing depths are very similar between the two studies, as demonstrated later on.
In the following, the 1D permafrost simulations for the location of the repository are described first,
followed by the 2D simulations across the Forsmark site.
1D reconstruction of last glacial cycle permafrost conditions at the repository location

The 1D modelling experiments described below, conducted for the repository location, are from/SKB 2006a, Section 3.4/.
The 1D calculations were carried out in two steps; i) by using constant surface temperatures, and ii)
using surface temperatures from the reconstruction of last glacial cycle conditions. The first task was
to investigate the ground surface temperatures required to reach different permafrost depths. A setof constant mean annual ground surface temperatures were assumed and applied to a 1D subsurface
model based on present-day reference data for Forsmark (Table 3-8, Figure 3-51). Note that since the1D permafrost experiments were conducted, these input data have been updated for the Forsmark site.
Updated input data on bedrock and groundwater properties were used for the 2D modelling studies/Hartikainen et al. 2010/, described in the second half of the present section.
The intention with the constant surface temperature simulations was to find steady state permafrost
depths. The modelling, however, showed that no steady state could be reached in a time frame of
two glaciation cycles. The results from this modelling effort for the time period of 120,000 years
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are shown in Figure 3-52. The modelling indicates that reaching a permafrost depth of 450 m at
Forsmark, without a heat contribution from spent fuel, requires a mean annual surface temperature of−8°C for around 120 kyrs. With the heat contribution from the fuel, surface temperatures need to belower than −8°C. The thickness of the cryopeg varies from some metres to over 70 m (Figure 3-52).
The results for permafrost development rates (aggradation rates) are shown in Figure 3-53.
In the 1D modelling of surface conditions, the objective was to define a correlation between airand surface temperatures in consideration of the insulating effects of surface covers. The so-calledn-factor concept /Lunardini 1978/ yields an empirical relationship between the mean annual surface
temperature Ts and the mean annual air temperature Ta as
Ts = f (Ta , Aa ; nfr , nth ) 							Equation 3-11

where Aa is the amplitude of annual air temperature variations, which can be defined as the half ofthe temperature difference between the monthly mean summer maximum and monthly mean winterTable 3-8. Site-specific reference physical properties and conditions of rock mass and ground­waterat Forsmark used for the 1D permafrost modelling /SKB 2006a/, mainly from site description model
Forsmark 1.2 /SKB 2005a/. Note that since the 1D permafrost experiments were conducted, thesedata have been updated. Updated input data on bedrock and ground­water properties were used forthe 2D modelling studies /Hartikainen et al. 2010/, described in the second half of the present section.
Rock mass

Thermal properties
Parameter

Unit

Reference

Min

Max

Temperature dependency

Thermal conductivity

W/(m·K)

3.6

3.2

4.0

11.0% decrease/100°C increase

Heat capacity

MJ/(m3·K)

2.3

2.1

2.5

27.5% increase/100°C increase

Thermal diffusivity

mm2/s

1.57

1.28

1.9

38.6% decrease/100°C increase

Reference

Min

Max

Reference

Min

Max

2

Interior Earth conditions by depth
0m

10,000 m

Parameter

Unit

Geothermal heat flow

mW/m

59.0

48.0

65.0

43.2

41.7

46.0

Radiogenic heat prod.

µW/m3

2.5

1.0

3.0

0.92

0.37

1.1

Reference temperature conditions by depth
Parameter

Unit

400 m

500 m

600 m

Temperature

°C

10.6

11.7

12.8

Temperature gradient

°C/km

11.0

Reference hydro-mechanical properties by depth
Parameter

Unit

100 m

200 m

300 m

≥ 400 mBulk density

kg/mYoung’s modulus

GPa

2,700

2,700

2,700

2,700

74

74

74

Poisson’s ratio

74

–

0.26

0.26

0.26

0.26

Kinematic porosity–

0.00001

0.00001

0.00001

0.00001

Total porosity–

0.00010.00010.00010.0001

Permeabilitym2

1.0·10–13

1.0·10–15

1.0·10–17

1.0·10–18

3Groundwater

Reference salinity concentrations by depth
Ionic solvent

Unit

10 m

100 m

300 m

500 m

≥ 1,000 m

2,000

Na+

g/m3

65

500

2,000

2,000

Ca2+

g/m3

60

250

1,200

1,000

3,500

Cl–

g/m3

15

1,300

5,500

5,500

10,000

SO42–

g/m3

20

170

500

400

50

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0Depth (km)

2

4

6

8

10

0

50

100

150

Temperature (°C)

200

250

Reference subsurface conditions
Reference subsurface conditions except minimum thermal conductivity
Reference subsurface conditions except maximum thermal conductivity
Reference subsurface conditions except maximum geothermal heat flow
Reference subsurface conditions except minimum geothermal heat flow

Figure 3-51. Modelled present-day bedrock temperatures at the Forsmark repository location using thesite-specific reference subsurface conditions in Table 3-8.
0
100

−2

200

−4

300
400

−8

500
−10
600

−12

700

−14

800

−16

−2

200

−4

600

−20

1000

60

80

100

120

−16

800

1000

Time (kyrs)

−12

700

900

40

−8

500

−18

20

−6

400

900

0

Surface temperature (°C)

100

300

−6

Depth (m)Depth (m)0Surface temperature (°C)

−20
0

20

40

60

80

100

120

Time (kyrs)

Figure 3-52. Left: Evolution of permafrost depth (thick lines) and related cryopegs (shadowed areas)
for constant ground surface temperatures of –2, –4, …, –20°C at the Forsmark repository location whenreference subsurface properties are assigned. Right: Evolution of permafrost depth for different constantground surface temperatures with and without the heat contribution from the repository (6,000 canisters).
Bold lines show calculations with heat from the repository, whereas thin lines show calculations with norepository heat contribution.

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Permafrost development rate (cm/a)

1000

100

10

1

−20
−12
−8
−6
−4

0.1

0.01

−2

0.001

10

Surface temperature (°C)
100

1000

10,000Time (yrs)

100,000

Figure 3-53. Evolution of permafrost development rate for constant ground surface temperatures of –2, –4,–6, –8, –12 and –20°C at Forsmark when site-specific subsurface properties are assigned.

minimum. nfr and nth are so-called freezing and thawing n-factors. These factors are ratios between
the surface and air freezing indexes, i.e. annual time integrals of associated temperatures below
the freezing point, and the surface and air thawing indexes, i.e. annual time integrals of associatedtemperatures above the freezing point, respectively. Depending on the type of surface cover, freezing
n-factors can vary between 0.1 and 1 whereas thawing n-factors can have values in the range from
0.3 to 1.5 /Lunardini 1978, Shur and Slavin-Borovskiy 1993, Taylor 1995, 2001, Klene et al. 2001,
Karunaratne and Burn 2004/. Figure 3-54 shows the relation of the mean annual air temperature
to the mean annual ground surface temperature at Forsmark obtained for a reference surface coverwhen vegetation and snow are considered and for a bare surface when no vegetation and snow exist.
The relationship is based on the assumption that ground surface and air temperature is constant over
the whole glacial cycle, when the surface is not covered by ice-sheet or submerged.
Since the factors of major importance to the subsurface conditions are the bedrock thermal properties
and the conditions affecting the geothermal heat flow, the development of permafrost was in /SKB 2006a/
investigated for extreme thermal diffusivities and geothermal heat flows as shown in Tables 3-8.
These results are not repeated here, since similar sets of sensitivity tests were made with updated
information on the geothermal properties at Forsmark with the 2D permafrost model /Hartikainen
et al. 2010/, see further down in this section.
After the sensitivity analysis using constant surface temperatures, permafrost development wasmodelled for time-varying site-specific surface conditions. For this purpose, the ice sheet model in
Section 3.1.4 was used to extract estimates of site-specific time series of ground-level annual air
temperatures and basal ice sheet temperatures, as well as information on site-specific ice thickness
variations. A description of these reconstructed temperatures for the last glacial cycle, including a
discussion and estimates of their uncertainties, is found in Appendix 1. The GIA modelling describedin Section 3.3.4 was used to produce input data on shore-level displacement.
The 1D permafrost model was used to reconstruct the permafrost conditions at the repository location
for the reconstruction of the last glacial cycle. It was also used for studying a case with extreme surface
conditions favouring permafrost growth /SKB 2006a/. The latter case is equivalent to the severe
permafrost case analyzed with the 2D model (reported in /Hartikainen et al. 2010/ and later in
Section 5.5.3), and is therefore not included in the present report. For the reconstruction of permafrost for the last glacial cycle, the air temperature curve seen in Figure 3-55 was applied. When thesite is covered by the reconstructed the Weichselian ice sheet, surface temperatures equal the basal
temperatures from the ice sheet model. When the site is submerged, as seen in the reconstructionof Weichselian shore-level variations, the surface temperature is set to +4°C.
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Mean annual ground surface temperature (°C)

15Reference vegetation and snow cover(nfr = 0.5, nth = 1.0)

10

No vegetation and snow cover(nfr = 1.0, nth = 1.3)

5

0

−5

−10

−15
−15

−10

−5

0

5

10

15

Mean annual air temperature (°C)

Figure 3-54. Relationship between the mean annual air temperature and the mean annual ground surface
temperature at Forsmark considering reference vegetation and snow cover, and no vegetation and snow cover.
15

Air temperature with ice and submerged conditions
Air temperature without ice and submerged conditions

10

0

-5

Air temperature (°C)

5

-10

-15
120

100

80

60

Time (kyrs before present)

40

20

0

Figure 3-55. Reconstructed air temperature curve for the Forsmark region for the past 120 kyrs, including
estimated submerged periods of the Eemian, Mid-Weichselian and Holocene interglacials (blue line), see
Appendix 1. The temperature scale shows absolute temperatures. The red line shows reconstructed lastglacial cycle air temperatures without a presence of an ice sheet and submerged periods. The blue line
includes periods of ice sheet coverage, i.e. it shows simulated basal ice temperatures for glaciated periods,
air temperatures for ice-free periods, and also a temperature of +4°C for submerged periods. The curves havebeen used as input to the SR-Site 1D and 2D permafrost simulations /SKB 2006a, Hartikainen et al. 2010/.
A submerged period during the Eemian was not included in the 1D permafrost simulations. The uncertainties
of the temperature curves are discussed in Appendix 1, and the implications for modelled permafrost depths
in the part of Section 3.4.4 describing the sensitivity studies made with the 2D permafrost model.

The results of the development of permafrost, perennially frozen subsurface and cryopeg for reference surface conditions (Table 3-8) with vegetation and snow cover, and considering heat from thespent fuel are shown in Figure 3-56. The uncertainty in the modelled permafrost and freezing depths
are analysed and discussed in the 2D permafrost modelling section below. The associated rates ofpermafrost aggradation and degradation are shown in Figure 3-57. The rather thick cryopeg layers
during glaciations are due to large water pressures created by warm-based ice sheets keeping the
subglacial bed unfrozen down to a temperature about –2°C.

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115

0

50

Depth (m)

100

150

200

permafrost line
perennially frozen
cryopeg

250

300
120

100

80

60

Time before present (kyrs)

40

0

20

Figure 3-56. Evolution of permafrost, perennially frozen subsurface and cryopeg at the repository locationin Forsmark considering reference surface conditions, reference subsurface properties and heat from thespent fuel, see also /SKB 2006a/. The uncertainty in the modelled permafrost and freezing depths areanalysed and discussed in the 2D permafrost modelling section below.

The results show that surface conditions can be seen as the driving force for the development of
permafrost, with the subsurface conditions and the heat from the spent fuel acting as either reducing
or enhancing factors. The results indicate that permafrost (0°C isotherm) can develop to near repository depth under the reference surface conditions when the insulating effects of vegetation and snow
cover are excluded.
Large water bodies will affect permafrost depth and will be associated with the presence of taliks. A talik
can exist beneath a water body when the water bottom temperature remains above the freezing point.
Based on field tests on temperature measurements in lakes e.g. /Yershov 1998, Burn 2002/, it wasdetermined that a two metres deep water body maintains its bottom temperature above 0°C and that awater body deeper than 8 m has a bottom temperature around +4°C. The size of the water body required
to retain a talik was investigated by permafrost modelling. To this end, permafrost development was
simulated in the vicinity of circular lakes with constant positive lake bottom temperatures and constantnegative lake bottom level bedrock temperatures. The results shown in Figure 3-58 (left panel) indicatethat an open talik, completely penetrating the permafrost, can survive beneath a circular shallow lake,
if its radius is greater than the thickness of surrounding undisturbed permafrost thickness. Furthermore,
the results in Figure 3-58 (right panel) demonstrate that a radius greater than 0.6 times the thickness ofsurrounding undisturbed permafrost is big enough for a deep circular lake to maintain an open talik.
The results apply for different lake bottom level bedrock temperatures.
For further sensitivity experiments conducted with the 1D permafrost model, see /SKB 2006a/.
In summary, the 1D permafrost study shows that the surface conditions are the driving force forthe development of permafrost at Forsmark, while the subsurface conditions and the heat from thespent fuel acted as either reducing or enhancing factors. Permafrost can aggregate from some cm tosome dm in a year whereas its degradation can take place several times faster, especially when thesurface temperature is increased above 0°C and permafrost degrades simultaneously from the bottom
upwards and from the top downwards. The results further show that the maximum permafrost depth
in the reconstruction of last glacial cycle conditions occurs at 70 kyrs BP, prior to the first main ice
sheet advance over the site. The permafrost depth is ~250 m at the repository location at this time.
However, the uncertainty of surface and subsurface conditions also needs to be taken into account,something that was done in /SKB 2006a/. In the present report, this is done in the following 2D
permafrost study.
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TR-10-49

6

1

5

0.8

4

0.6

3

0.4

2

0.2

1

0
120

100

80

60

Time (kyrs BP)

40

20

0

Permafrost degradation rate (m/a)

Permafrost aggradation rate (m/a)

1.2

0Figure 3-57. Evolution of permafrost aggradation and degradation rates at the repository location in
Forsmark considering reference surface conditions with vegetation and snow cover, reference subsurface
properties and heat from the spent fuel, see /SKB 2006a/.
lake radius 420 m, bottom temp +0.1°C+0.1 °C0lake radius 250 m, bottom temp +4°C

−8 °C

+4 °C

0

420 m

−8 °C

250 m

Permafrost

100

Time (ka)
~1

200

~4

300

~10
~25

400

~120

Ground with T ≥ 0 °C
500

0

200

400

600

800

1000

Distance from centre of lake (m)Depth below lake bottom (m)Depth below lake bottom (m)Permafrost
100

Time (ka)
~1

200

~4

300

~10
~25

400

~120

Ground with T ≥ 0 °C
500

0

200

400

600

800

1000

Distance from centre of lake (m)Figure 3-58. Evolution of permafrost in the vicinity of a circular lake (red colour) at Forsmark. Left panel:
the lake radius is 420 m and the constant lake bottom temperature +0.1°C. Right panel: the lake radius is
250 m and the constant lake bottom temperature +4°C. In both cases, a constant bedrock temperature of–8°C prevail at the lake bottom level.

2D reconstruction of last glacial cycle permafrost conditions at the Forsmark site

The 2D permafrost study is a continuation and complement to the 1D study performed for SR-Can/SKB 2006a, Section 3.4 and 4.4.1/. The overall aim is to exemplify how permafrost and frozen
ground develops over the Forsmark site (and not only at the repository location) during cold but
ice-free periglacial phases of glacial cycles, and to further demonstrate the effects of uncertainties
associated with processes and phenomena of importance for permafrost development. To this end, the2D study provides a numerical estimation of the development of permafrost and perennially frozen
ground along a profile covering a major portion of the Forsmark site and crossing the repository location (Figure 3-59), using site-specific surface and subsurface conditions. The objective is to investigate
effects of multidimensional features of surface and subsurface conditions on the occurrence, development and distribution of permafrost and perennially frozen ground. Special emphasis is put on themodelling of surface conditions, including climate (temperature and humidity), soil, vegetation, water
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117

1635000

1640000

1645000

6707000
6702000
6697000

2D profile studied by permafrost modelling

0

1

2

Suggested location for repository for spent nuclear fuel6692000

6692000

6697000

6702000

6707000

6712000

1630000

6712000

1625000

3 km
Svensk Kärnbränslehantering AB

Forsmark site (biosphere model domain)© Lantmäteriverket Gävle 2007.
Consent I 2007/1092.
Coordinate system: RT90 2.5 gon W

Future lakes
1625000

1630000

1635000

1640000

1645000

Figure 3-59. The ~15 km long profile analysed in the 2D permafrost modelling study (red line), repository
location (grey), and the Forsmark site (stippled) as defined and analysed in the SR-Site biosphere
programme, e.g. /SKB 2010d, e/.

bodies and topography, lateral variations in bedrock and surface physical properties and heat generationfrom the spent fuel. In addition, groundwater flow and salt transport processes are considered (in 2D);
processes that were neglected in the 1D SR-Can study.
The mathematical model for freezing and thawing of the subsurface, based on the approach describedin /Hartikainen and Mikkola 2006/, has been used in /Hartikainen and Mikkola 2004, Hartikainen
2006, SKB 2006a/. The model is capable of describing heat and mass transfer in a porous medium
with freezing of groundwater depending on groundwater pressure and salt concentration and freezinginduced groundwater flow. In the present 2D study, by using the theory of mixtures and basic principles
of continuum mechanics and thermodynamics, the mathematical model is reformulated in terms ofGibbs free energy functions by taking into account the effects of temperature, pressure and salinity
concentration on the properties of groundwater and ice in a more accurate way. The reformulated

118

TR-10-49thermo-hydro-chemically coupled model consists of balance laws of mass, momentum and energy, andthe constitutive laws of state and diffusion. A new feature of the model is the capability of describing
the exclusion of salt during freezing and the density dependent groundwater flow in unfrozen andpartially frozen ground. In addition, the effects of pressure and salinity concentration on the freezing
temperature of groundwater are considered more accurately. A more detailed description of the 2D
permafrost model used for the reconstruction of last glacial cycle permafrost conditions is found in/­Hartikainen et al. 2010/.
The same climate development as used for the 1D permafrost simulations was used, complemented
with more comprehensive sensitivity studies on e.g. air temperature uncertainties. However, the 2D
reconstruction of last glacial cycle permafrost conditions do not consider the entire last glacial cycle,
but instead only the ice-free period from 115 to 70 kyrs BP, i.e. the period prior to the first phase of ice
sheet coverage in the reconstruction of the Weichselian ice sheet (see Section 3.1.4). This is motivatedby the fact that this period ends with the deepest permafrost reconstructed at the site for the entire lastglacial cycle, see /SKB 2006a/ and description of this 1D permafrost study above. However, for thesevere permafrost case (Section 5.5), 2D permafrost simulations are carried out for the full glacial
cycle, from 115 kyrs BP to present. Initial conditions are assumed to correspond to the present-day
conditions (i.e. bedrock temperature, groundwater, salinity and pressure). Heat generated by the spent
fuel in the repository is included in the simulations at 112 kyrs BP (see also Section 3.4.7).
Based on the results of the permafrost simulations in /SKB 2006a, Section 3.4 and 4.4.1/, the present
analysis focuses on the following issues:
Subsurface conditions:
•
•
•
•
•

Spatial variation of bedrock and soil thermal and hydraulic properties.
Site-specific geothermal heat flow.
Convectional heat transfer due to 2D groundwater flow.
Heat generation from the repository for spent fuel.
Salinity exclusion due to freezing and salt transport.

Surface conditions:••••

Air temperature and its variation over time.
Topographical features including slope and relief.
Water bodies (sea and lakes), including the effect of isostatic uplift.
Vegetation and snow cover as related to climate and topographical conditions.

The profile is located along the, very low, regional topographic gradient. In the direction of the profile,
it covers a major part of the Forsmark site as defined by the biosphere model domain (Figure 3-59).
For further description of the selection of the profile location, see /Hartikainen et al. 2010 Appendix A/.
Subsurface properties and conditions
The model domain encompasses an approximately 15 km long and 10 km deep vertical section
consisting of six soil layers, 23 rock domains and 31 deformation zones as shown in Figures 3-59
and 3-60. Figure 3-61 illustrates the thermally different rock mass domains, the deformation zonesof the upper 2.1 km of the model domain, as well as the present day Baltic Sea-level. The soil
layers are shown in Figure 3-62. The thermal and hydraulic properties of soil domains are found in/Hartikainen et al. 2010, Table 2-1/, whereas Table 2-2 and Table 2-3 in the same publication give thethermal and hydraulic properties of the rock mass domains. Hydraulic properties of the deformation
zones are presented in /Hartikainen et al. 2010 Appendix D/. The thermal properties of deformation
zones are assumed identical to those of the corresponding rock mass domain. The bedrock below thedepth of 2.1 km is assumed to have the same thermal and hydraulic properties as the rock domain
in which the repository is located. The material properties of the bedrock are described in detailin /Hartikainen et al. 2010, Appendices D and E/. For the treatment of e.g. unfrozen groundwatercontent, also see /Hartikainen et al. 2010/. The numerical simulations were carried out using an
unstructured finite element mesh of linear triangle elements. The mean grid spacing varied fromless than 10 m close to the ground surface to about 300 m at the bottom of the model domain. The
maximum time step for the adaptive time-integration scheme was limited to 100 years.
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119

Figure 3-60. The model domain used in the 2D permafrost simulations. Colours from blue to red signify
the 23 rock domains considered, and grey the 31 deformation zones. Physical and thermal characteristics
for the rock domains and deformations zones were obtained from the site investigation programme at
Forsmark, see /Hartikainen et al. 2010/. The six soil layers described in /Hartikainen 2010 Appendix B/ andFigure 3-62, are too thin to be seen in this figure. The x-axis is directed eastward and y-axis northward.
The RT-90 coordinates of the south-west side are x = 1,628,228 m, y = 6,696,472 m.

Figure 3-61. The upper 2.1 km of the 2D model domain. Colours from blue to red signify rock domains
with different thermal diffusivity. Numbers from 1 to 10 indicate the rock domains of Table 2-2 in/Hartikainen et al. 2010/. The location of deformation zones is illustrated in grey and the repository inwhite, at a depth of 450 m starting at a distance of about 4,000 m along the profile. The thin blue layeron the top surface represents the present day Baltic Sea-level. Further description and references for thesedata are found in /Hartikainen 2010 Appendix C/.

120

TR-10-49

Figure 3-62. The soil layers of the 2D model domain. Grey colour represents bedrock and blue the present
day Baltic Sea. The gyttja, peat and sand layers are too thin to be seen in this figure. In addition, because thefirst 2,000 m of the model domain is outside the area of detailed site investigations, no information about thesoil cover is used here. Further description and references for these data are found in /Hartikainen et al. 2010
Appendix B/. Note that the vertical axis has a scale which gives a strong vertical exaggeration of topography.

The geothermal heat flow is an important subsurface parameter for permafrost modelling. The geo­
thermal heat flow, crustal radiogenic heat production and ground temperature depth to ~1 km arebased on /Sundberg et al. 2009/, see also /Hartikainen et al. 2010, Appendices B–F/. For a descriptionof how these parameters are modelled for depths from ~1 km to 10 km, see /Hartikainen et al. 2010/.
Figure 3-63 shows the initial ground temperature calculated for mean thermal properties.
The initial salinity concentration of groundwater for depths to ~1.5 km is described in /Hartikainen
et al. 2010 Appendix F/, and for depths from ~1.5 km to 10 km it is obtained as a stationary solution
using present concentration values on the surface and at the depth of 1.5 km. Figure 3-64 shows theinitial salinity distribution in the model domain.

Figure 3-63. Distribution of initial ground temperature for mean thermal properties used in the 2D
permafrost modelling. For a description of these properties, see /Hartikainen et al. 2010, Table 2-2/.

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121

Figure 3-64. Distribution of initial salinity concentration of groundwater /Hartikainen et al. 2010/.

The present day groundwater pressure is determined as hydrostatic pressure based on the initial conditions of ground temperature and groundwater salinity concentration as well as the hydrostatic pressure
at the bed of the Baltic Sea. The results for initial groundwater pressure are shown in Figure 3-65.
For a description of the localisation of the repository along the profile and the treatment of heatproduced by the spent nuclear fuel, see /Hartikainen et al. 2010/.
Surface properties and conditions
Due to significant uncertainties associated with descriptions of the surface conditions, two variants,
one humid and one dry, are analyzed. This is expected to yield a lower and upper limit for thepermafrost and perennially frozen ground development, see further below.
As mentioned above, the 2D reconstruction of last glacial cycle permafrost conditions analyse theperiod from 115 to 70 kyrs BP, that is a ~85 kyr long period from the Eemian interglacial up to thefirst phase of glacial conditions as reconstructed for the Forsmark site /SKB 2006a/. Based on the
Köppen climate classification /Lohmann et al. 1993/ and long-term observations /Eugster et al. 2000/
both climate variants may be divided into three climate zones: Boreal, Subarctic and Arctic, which
may be characterized by the annual mean air temperature, the monthly mean maximum summer air
temperature and the monthly mean minimum winter air temperature, together with the monthly mean
maximum summer and winter precipitations (Table 3-11). The large ranges in the monthly mean
air temperatures and precipitations are explained by the fact that the description and modelling of
surface conditions involve significant uncertainties, see Appendix 1 and later in Section 3.4.
The shore-level development (Section 4.4.2) and evolution of lakes along the profile, as described in
the SR-Site biosphere studies, are included in the modelling, see /Hartikainen et al. 2010 Appendix I/.
As a result of isostatic uplift two lakes are formed along the profile, in the area today covered by theBaltic Sea (Figure 3-59).
The vegetation and snow cover are closely related to the prevailing climate conditions as well as
to other landscape parameters. The presence of different vegetation types within different biomes
has been described in numerous publications, e.g. /Breckle 2002, Archibold 1994/ and the regional
response of vegetation types to e.g. periglacial climate conditions has been studied by /Kjellström
et al. 2009b, Kjellström et al. 2010a/. In the present study, a detailed picture of the distribution ofprevailing major vegetation types for Boreal, Subarctic and Arctic climate zones for the reconstruction of last glacial cycle conditions was described by using local variations in TWI (Table 3-9).

122

TR-10-49

Figure 3-65. Distribution of initial groundwater pressure associated with the temperature and salinity
concentration shown in Figure 3-63 and 3-64.
Table 3-9. Main vegetation cover types at Forsmark for three different climate zones and corresponding Topographic Wetness Index (TWI) values. For a more detailed description of the useof e.g. the TWI index in this study, see /Hartikainen et al. 2010 Appendix J/.
Climatezone

Vegetation cover type and TWI values
ForestShrubsGrassland

BarrenPeatlandBoreal

Mixed

–

–

–

Peat

Moderate/short

Short

–

Peat

TWI > 0

TWI > 13.2

Subarctic

Tree line/
tundra forest
TWI > 13.2

10.9 ≤ TWI ≤ 13.2

TWI < 10.9

Arctic–Short

Tussocks

Bare

Peat

TWI > 13.2

10.9 ≤ TWI ≤ 13.2

TWI < 10.9

TWI > 13.2

TWI > 13.2

Table 3-10. Freezing (winter) and thawing (summer) n-factors for the reconstruction of last glacial
cycle conditions. See also /Hartikainen et al. 2010/. The numbers within brackets in the last fourcolumns are Topographic Wetness Index (TWI) intervals.
Climaticzone

SeasonMonthly mean
max./min.
air temperature(°C)

Monthly meanmax.
precipitation(mm/month)n-factors for specified surface conditions andassociated TWI intervals
Dry
Fresh-moist
Wet
Peatland(< 10.9)(10.9–13.2)(> 13.2)(> 13.2)

Borealsummer

+10 – +20

10–60

1.2–1.4

1.2–1.4

1.2–1.4

1.0

winter–15 – +5

1–50

0.2–0.3

0.2–0.3

0.2–0.3

0.1

summer+5 – +15

5–60

0.8–1.0

0.7–0.9

0.6–0.8

0.8–1.0

winter

–25 – –10

1–40

0.5–0.9

0.4–0.8

0.2–0.6

0.1–0.4

summer< +5 – +10

5–60

1.0–1.2

0.5–0.8

0.4–0.7

0.6–0.9

winter

–35 – –20

1–40

0.7–1.0

0.4–0.8

0.3–0.6

0.2–0.5

Subarctic
ArcticTR-10-49

123

Table 3-11. Climate information for the reconstruction of last glacial cycle conditions.
Climate zone

Annual mean
air temperature(°C)

Monthly mean
max. summer air
temperature (°C)

Monthly mean
min. winter air
temperature (°C)

Monthly mean
max. summer
precipitation(mm/monthMonthly mean
max. winterprecipitation(mm/month)

Boreal>0+10 – +20

−15 – +5

10–60

1–50

Subarctic

0 – –6+5 – +15−25 – –10

5–60

1–40

Arctic< –6< +5 – +10−35 – –20

5–60

1–40

The spatial resolution of the TWI information was 20·20 m. For a more detailed description of the
adopted TWI approach, see /Hartikainen et al. 2010/.
The impact of vegetation, snow cover and other climate factors on the ground surface temperatureand permafrost development is well investigated /Washburn 1979, Williams and Smith 1989, Yershov
1998, French 2007/. In general, an annual mean air temperature ranging between –9 and –1°C is
required to build up permafrost for a majority of surface covers, depending on other climatic conditions and topography. The surface conditions can be modelled by thermodynamic and hydrodynamic
models using surface energy and water balance equations and information on climate conditions
and topography such as radiation, precipitation, cloudiness and wind, and their annual and diurnal
variation /Riseborough et al. 2008/. However, these models are unsuitable for the long time spans
associated with glacial cycles, since no climate data besides the air temperature can be adequately
constructed. Therefore, an empirical approach based on n-factors /Lunardini 1978/, i.e. statisticalcorrelations between air and ground surface temperature, has been used to construct the groundsurface temperature from the reconstructed air temperature (Table 3-10). For a detailed motivation
for and description of the n-factors approach as an approach to treatment of surface conditions, see/­Hartikainen et al. 2010/.
When the site is submerged by the Baltic Sea or the two lakes, the ground surface temperature was
set to +4°C for water depths greater than 6 m. For shallower depths the temperature was interpolated
between the value of 4°C and the prevailing ground surface temperature.
Figure 3-66 exemplifies the evolution of modelled ground surface temperatures for the reconstruction of last glacial cycle conditions at one location along the profile with fresh-moist surface condi14

Ground surface temperature

12

humid variant
dry variant
range

10
8

Temperature (°C)

6
4
2
0
−2
−4
−6
−8
−10
−12
−14
115

Air temperature← 111.5
110

105

100

95

90

← 74
85

80

75

70

Time before present (kyrs)

Figure 3-66. Example of evolution of air temperature and modelled ground surface temperatures for a partof the profile with fresh-moist surface conditions (~4,800 m from the south-west side of the profile, above
the repository, see Figure 3-59) for the reconstruction of last glacial cycle conditions.

124

TR-10-49

tions (at ~4,800 m from the south-west side). Examples of modelled ground surface temperatures
for other surface condition types (wet, dry and peatland) are found in /Hartikainen et al. 2010/.
Figure 3-67 illustrates the modelled ground surface temperatures along the profile at four different
times during the last glacial cycle, considering different climate conditions: 111.5 kyrs BP – sub­arctic and partially submerged, 103.5 kyrs BP – boreal, 95 kyrs BP – subarctic, 70 kyrs BP – arctic.
The results concerning the humid climate variant are obtained using the lowest n-factors of Table 2-7
and 2-8 in /Hartikainen et al. 2010/, whereas the highest n-factors from the same tables yield theresults for the dry climate variant.
The ground surface temperature of +4°C indicates the times when the profile is submerged by theBaltic Sea or the two lakes. The dry surface condition type generates the lowest ground surface
temperatures, and the peatland condition the highest ground surface temperatures in the subarctic
and arctic climate zones. Furthermore, due to thin vegetation and snow cover, the dry variant of thereconstruction of last glacial cycle conditions results in the lowest ground surface temperatures inthe subarctic and arctic climate zones but highest in the boreal climate zone.
14103.5 ka BP - Boreal climate
111.5 ka BP - Subarctic climate, partly submerged
95 ka BP - Subarctic climate
70 ka BP - Arctic climate

12

10

Mean annual ground surface temperature (°C)

8

6

4

2

0

−2

−4

−6

−8

−10

3.55 km
−12

0

1

2

3

4.783 km4

5

6

11.6 km
7

8

9

10

11

14.6 km
12

13

14

15

Distance from south−western side of profile (km)Figure 3-67. Modelled ground surface temperatures along the profile at four different times for the lastglacial cycle case considering different climate conditions: Boreal climate at 103.5 kyrs BP, Subarctic climate
and partially submerged conditions at 111.5 kyrs BP, Subarctic climate at 95 kyrs BP, Arctic climate at
70 kyrs BP. The solid border lines of the shadowed areas indicate the dry variant of the climate case anddashed ones the humid variant of the reconstruction of last glacial cycle conditions. The dashed verticalline denotes the location used as example for the fresh-moist surface condition types (Figure 3-66).

TR-10-49

125

Descriptions of cases simulated with the 2D permafrost model
For the reconstruction of permafrost conditions for the last glacial cycle, two main simulations weremade constituting one dry and one humid climate variant. The minimum values of n-factors shownin Table 3-10 describe the humid variant whereas their maximum values yield the dry variant. Mean
subsurface properties and conditions are used for both variants (e.g. mean thermal and mechanicalproperties of soil and bedrock, and chemical properties of groundwater, see /Hartikainen et al. 2010/).
Main simulations:
1. Dry climate variant of last glacial cycle.
2. Humid climate variant of last glacial cycle.
The uncertainties related to surface conditions including moisture conditions, vegetation and snow coveras well as ground surface temperature are considered in the main simulations by means of the humid
and dry variant of the climate cases studied. Variations in hydraulic properties and ionic composition ofgroundwater are considered to be of negligible significance in regard to permafrost development, andhence were not included in the sensitivity analyses. Following this approach, and taking results fromprevious studies into account /SKB 2006a/, the following sensitivity studies were done:
1. uncertainty in air temperature,
2. uncertainty in geothermal heat flow,
3. uncertainty in bedrock thermal conductivity,
4. uncertainty in bedrock thermal diffusivity,
5. combination of uncertainties in bedrock thermal properties,
6. combination of uncertainties in surface and bedrock thermal properties,
7. combination of uncertainties in air temperature, surface conditions and bedrock thermal properties,
8. influence of heat from the repository,
9. convective heat transfer by groundwater flow.
Description and results of main simulations

Resulting ground temperatures and extent of perennially frozen ground at times of 111.5, 95, 74,
70 kyrs BP are illustrated in Figure 3-68 and 3-69 for the humid and dry variant of the reconstructionof last glacial cycle conditions. The selected times in the figures represent the four different situations regarding climate zone, shore-level development and permafrost conditions:•111.5 kyrs BP: subarctic climate prevails, the profile is partially submerged by the Baltic and theexposed ground surface is partially underlain by permafrost. The temperature is at a maximum
within the repository,•

95 kyrs BP: subarctic climate prevails and discontinuous permafrost is developing,•

74 kyrs BP: subarctic climate is turning into a warmer boreal climate and permafrost is degrading,changing from continuous or sporadic permafrost,•

70 kyrs BP: arctic climate prevails and continuous permafrost reaches its maximum depth.

The evolution of maximum permafrost depth, maximum depth of perennially frozen ground and the
maximum depths of the −2 and –4°C isotherms over the repository for the reconstruction of last glacial
cycle conditions (both dry and humid climate variants) is shown in Figure 3-70. Figure 3-71 shows thesame results for the whole profile and for the repository location, as well as the spatial extent of permafrost summarised along the investigated profile. The −2°C temperature corresponds to the temperature
criterion used for freezing of the back-fill material in the deposition tunnels, while –4°C constitute thetemperature criterion used in the safety assessment for freezing of the buffer clay /SKB 2011/.
The maximum permafrost depths, maximum depths of perennially frozen ground, and the extent
of permafrost distribution are summarised in Tables 3-12 and 3-13. In addition, Table 3-14 andTable 3-15 show the evolution of permafrost depth and depth of perennially frozen ground over
the repository and whole profile, respectively.

126

TR-10-49

111.5 kyrs BP

95 kyrs BP

74 kyrs BP

70 kyrs BP

Figure 3-68. Temperature contours in (°C) and the extent of perennially frozen ground (light colour) andpermafrost (0°C isotherm) at times 111.5, 95, 74, 70 kyrs BP for the humid variant of the reconstruction oflast glacial cycle conditions. Blue colour on the top of the profile at 111.5 BP kyrs shows the Baltic Sea.
The yellow rectangle indicates the location of the repository.

TR-10-49

127

111.5 kyrs BP

95 kyrs BP

74 kyrs BP

70 kyrs BP

Figure 3-69. Temperature contours in (°C) and the extent of perennially frozen ground (light colour) andpermafrost (0°C isotherm) at times 111.5, 95, 74, 70 kyrs BP for the dry variant of the reconstruction of lastglacial cycle conditions. Blue colour on the top of the profile at 111.5 kyrs BP shows the Baltic Sea. The
yellow rectangle indicates the location of the repository.

128

TR-10-49050

Depth below ground surface (m)

100

150

200

250

Maximum permafrost depth
300

350

400

humid variant
dry variant

Maximum depth of perennially frozen groundhumid variant
dry variant

Maximum depth of −2 °C and −4 °C isotherm

450−2 °C,−2 °C,

500
115

110

−4 °C, humid variant−4 °C, dry variant

105

100

95

90

85

Time before present (kyrs)

80

75

70

Figure 3-70. Evolution of maximum permafrost depth, maximum depth of perennially frozen ground andmaximum depths of the −2 and –4°C isotherms over the repository for the reconstruction of last glacial
cycle conditions. The upper permafrost surface, for periods of degradation from above, is not shown. The
shaded area in blue and red represents the range when considering the dry and humid climate variants.
The lilac colour indicates that the results for permafrost and perennially frozen ground overlap.

TR-10-49

129

Percent distribution along the entire profile

100
90
80
70
60
50
40
30

Permafrost extent

20

dry variant
humid
variant10050Depth below ground surface (m)

100

150

200

Maximum permafrost depth
250

300

humid variant
dry variant

Maximum depth of perennially frozen groundhumid variant
dry variant

350

Maximum depth of −2 °C and −4 °C isotherms
400−2 °C,−2 °C,

450
115

110

−4 °C, humid variant−4 °C, dry variant
105

100

95

90

Time before present (kyrs)

85

80

75

70

Figure 3-71. Evolution of maximum permafrost depth, maximum depth of perennially frozen ground andmaximum depth of −2 and –4°C isotherms over the whole profile for the reconstruction of last glacial cycle
conditions. The figure also shows the percent permafrost distribution along the profile. The transition
from sporadic to discontinuous permafrost occurs at 50% permafrost coverage and from discontinuous tocontinuous permafrost at 90% coverage. The upper permafrost surface, for periods of permafrost degradation from above, is not shown. The shaded area in blue and red represents the range when considering the
dry and humid climate variants. The lilac colour indicates that the results for permafrost and perenniallyfrozen ground overlap.

130TR-10-49

Table 3-12. Times of permafrost occurrence and associated maximum permafrost depth andmaximum depth of perennially frozen ground over the whole profile and over the repository forthe humid variant reconstruction of last glacial cycle conditions. The table also gives the timeand horizontal location of maximum permafrost depth and the percent extent of permafrost
distribution.

Maximum depth of perenniallyfrozen ground (m)Distance from the south-westside (m)Time of occurrence of maximum
permafrost depth (ka BP)

Maximum permafrost depth (m)Percent extent of permafrost
distribution (%)

Over the repository
Maximum depth of perenniallyfrozen ground (m)Distance from the south-westside (m)Time of occurrence of maximum
permafrost depth (ka BP)

Over the whole profile
Maximum permafrost depth (m)Time periodsof permafrost
occurrence(kyrs BP)

113–111.7

31

112

4,960

29

34

31

112

4,960

29

111.3–109.1

42

109.8

7,960

39

62

18

111

5,540

17

105.4–105.1

10

105.2

14,290

9

71

12

105.2

5,540

11

104.6–104.2

2

104.4

14,320

1

58

6

104.4

5,540

5

98–88.6

77

90.3

12,750

73

94

55

90.3

4,970

52

87.5–87.2

1

87.4

14,320

4

50

6

87.4

5,540

5

86.9–83.4

129

85.5

11,620

125

100

95

85.6

5,160

92

78.4–74188

74.4

11,630

182

100

152

74.4

5,120

148

72.6–70221

70

11,630

212

100

182

70.1

5,110

176

Table 3-13. Times of permafrost occurrence and associated maximum permafrost depth andmaximum depth of perennially frozen ground over the whole profile and over the repository forthe dry variant reconstruction of last glacial cycle conditions. The table also gives the time andhorizontal location of maximum permafrost depth and the percent extent of permafrost distribution.

Time of occurrence of maximum
permafrost depth (ka BP)

Distance from the south-westside (m)Maximum depth of perenniallyfrozen ground (m)

104

109.7

7,950

98

74

84

112

4,100

79

105.5–104

48

104.4

13,970

46

94

39

105.2

5,540

38

99.9–83.2

189

85.1

12,770

180

100

143

85.5

5,652

137

79.2–73.8

255

74.3

12,810

249

100

215

74.3

5,650

204

73.1–70301

7012,820

295

100

259

70

5,650

246

TR-10-49

Maximum permafrost depth (m)Percent extent of permafrost
distribution (%)

113.6–106.4

Distance from the south-westside (m)Maximum depth of perenniallyfrozen ground (m)Over the repository

Time of occurrence of maximum
permafrost depth (ka BP)

Over the whole profile
Maximum permafrost depth (m)Time periodsof permafrost
occurrence(kyrs BP)131

Table 3-14. Evolution of maximum permafrost depth and maximum depth of perennially frozen
ground over the repository modelled for the reconstruction of last glacial cycle conditions. The
table also shows the prevailing mean annual air temperatures.

Humid variant of the reconstruction of last glacial cycle

Dry variant of the reconstruction of last glacial cycle

Maximum depthof perennially frozen
ground (m)Dry variant of the reconstruction of last glacial cycle

Maximum permafrost
depth (m)Humid variant of the reconstruction of last glacial cycle

Time before Mean annual
present
air temperature
(kyr)
(°C)

112

–3.4

31

84

29

79

111.5

–0.6

0

3

0

1

110

–3.5

15

47

15

46
21

108

–1.5

0

22

0

106

2.5

0

0

0

0

104

–0.6

0

6

0

5

102

1.6

0

0

0

0

100

–0.3

0

0

0

0

98

–1.9

2

30

0

28

96

–2.9

18

62

16

60

95

–3.2

23

69

20

66

94

–3.9

35

84

33

81

92

–4.8

49

108

47

103

90

–4.1

45

112

43

107

88

–0.9

0

15

0

13

86

–6.4

73

116

71

113

84

–3.1

31

108

29

104

82

2.1

0

0

0

0

80

0.4

0

0

0

0
50

78

–2.5

12

52

10

76

–4.2

45

101

43

97

75

–10.4

135

186

131

179

74

–1.3

85

180

77

168

72

–3.6

33

86

31

82

70

–10

181

259

175

246

132

TR-10-49

Dry variant of the reconstruction of last glacial cycle

Maximum depth of
perennially frozen
ground (m)Humid variant of the reconstruction of last glacial cycle

Maximum permafrost
depth (m)Dry variant of the reconstruction of last glacial cycle

Time before Mean annual
present
air temperature
(kyr)
(°C)

Humid variant of the reconstruction of last glacial cycle

Table 3-15. Evolution of maximum permafrost depth and maximum depth of perennially frozen
ground over the whole profile modelled for the reconstruction of last glacial cycle conditions.
The table also shows the prevailing mean annual air temperatures.

112

–3.4

31

84

29

86

111.5

–0.6

0

12

0

10

110

–3.5

37

95

34

89

108

–1.5

0

41

0

38

106

2.5

0

0

0

0
20

104

–0.6

0

23

0

102

1.6

0

0

0

0

100

–0.3

0

0

0

0
32

98

–1.9

0

34

0

96

–2.9

29

93

26

89

95

–3.2

34

101

31

97

94

–3.9

49

119

46

114

92

–4.8

70

151

66

145
155

90

–4.1

70

162

65

88

–0.9

0

7

0

6

86

–6.4

96

153

93

147
147

84

–3.1

50

155

46

82

2.1

0

0

0

0

80

0.4

0

0

0

0

78

–2.5

1

59

1

56

76

–4.2

55

122

52

117

75

–10.4

165

222

160

212

74

–1.3

150

239

140

222

72

–3.6

44

110

41

105

70

–10

221

301

212

295

Figure 3-72 and 3-73 show vertical bedrock temperature profiles in the middle of the repository
location (with fresh-moist surface cover) for the humid and dry variants of the reconstruction of lastglacial cycle conditions. Similar examples of vertical bedrock temperature profiles, including othersurface cover types (wet, dry, peatland) are found in /Hartikainen et al. 2010/.
The modelling results for the reconstruction of last glacial cycle conditions show that annual mean
ground surface temperature at the Forsmark site can vary considerably in time and location dependingon the prevailing climate zone and surface moisture condition type, see Figures 3-66 and 3-67, and/Hartikainen et al. 2010/. For the first ~50 kyrs of the reconstructed last glacial cycle, the dry surface
condition resulted in arctic climate zone ground surface temperatures below −11°C and boreal climate
ground surface temperatures over +6°C. The wet surface condition yielded milder ground surface
temperatures in both the arctic and boreal climate zones ranging from below −3°C to over +5°C overthe same time period. Furthermore, the peatland surface condition, having the moistest surface conditions, resulted in even milder ground surface temperatures ranging approximately between −2°C and+4°C. In addition, the difference in ground surface temperature could be several degrees from place toTR-10-49

133

Temperature (°C)

Elevation above the present sea level (m)

−25
0

−20

−15

−10

−5

0

5

10

15

20

25

30

35

−100
−200
−300
−400

30.4 °C

Repository depth

8.6 °C

−500
−600
−700
−800
−900

At 115 kyrs BP (initial state)
At 111.5 kyrs BP (repository level max.)
At 103.5 kyrs BP (ground surface level max.)
At 95 kyrs BP (permafrost develops)
At 74 kyrs BP (permafrost degrades)
At 70.2 kyrs BP (ground surface level min.)
At 70 kyrs BP (repository level min.)−1000

Figure 3-72. Bedrock temperature along a vertical profile located in the middle of the repository at
4,783 m from the south-western side of the profile the for the humid variant of the reconstruction of lastglacial cycle conditions. The surface cover type reflects fresh-moist conditions. The white colour envelope
represents the range of bedrock temperature fluctuation, simulated over the time period of 115–70 kyrs BP.
For more examples of modelled bedrock temperatures, see /Hartikainen et al. 2010/.

Elevation above the present sea level (m)

−25
0

−20

−15

−10

−5

Temperature (°C)

0

5

10

15

20

25

30

35

−100
−200
−300
−400
−500

30.1 °C

Repository depth

7.0 °C

−600
−700
−800
−900

At 115 kyrs BP (initial state)
At 111.5 kyrs BP (repository level max.)
At 103.5 kyrs BP (ground surface level max.)
At 95 kyrs BP (permafrost develops)
At 74 kyrs BP (permafrost degrades)
At 70.2 kyrs BP (ground surface level min.)
At 70 kyrs BP (repository level min.)−1000

Figure 3-73. Ground temperature along a vertical profile located in the middle of the repository at 4,783 mfrom the south-western side of the profile for the dry variant of the reconstruction of last glacial cycle
conditions. The surface cover type reflects fresh-moist conditions. The white colour envelope represents
the range of bedrock temperature fluctuation, simulated over the time period of 115–70 kyrs BP. For more
examples of modelled bedrock temperatures, see /Hartikainen et al. 2010/.

134

TR-10-49

place along the investigated profile, see Figure 3-67. The same figure shows that the spatial fluctuationin ground surface temperature is largest in the arctic climate zone when the air temperature is lowest,
and lowest in the boreal climate zone when the air temperature is close to 0°C.
The simulation results (Figure 3-70) demonstrate that development of discontinuous permafrost
is characteristic for the humid climate variant of the reconstruction of last glacial cycle conditions
whereas continuous permafrost development prevails for long periods of the dry climate variant. In
all cases permafrost became a continuous spatial distribution when the annual mean air temperature
decreased below –6°C and the maximum permafrost depth exceeded a depth of ~50 m.
During some periods, short distinct changes in climate results in phases of fast decay and growthof permafrost, such as at around 74 kyrs BP (Figure 3-71). Similar rates of permafrost growth anddecay have been reported in /Delisle 1998, Kukkonen and Safanda 2001/.
The variation in ground surface temperature has the strongest effect on bedrock temperature fluctuations within the upper part of the vertical profile, see Figures 3-71 and 3-72 and /Hartikainen et al. 2010/.
The fluctuation is largest at the surface and reduces considerably with depth except where the heatfrom the repository has a large influence.
Regarding the evolution of permafrost depth and depth of perennially frozen ground, the results forthe reconstruction of the last glacial cycle that considered variations in surface moisture conditions
and mean bedrock thermal properties indicated that permafrost (defined by the 0°C isotherm) canreach a depth between ~180 m and ~260 m over the repository (Figure 3-70) and a depth between~220 m and ~300 m outside the repository (Figure 3-71) in a time frame of 45,000 years. The maximum depth of perennially frozen ground can range between ~180 m and ~250 m over the repository
and between ~210 m and ~300 m over the whole site at the same time. As seen from the 1D permafrost
simulations of the full last glacial cycle (Figure 3-56), these numbers represent the largest depths
during the entire glacial cycle.
A comparison between results from the present 2D modelling for the reconstruction of last glacial
cycle conditions and the corresponding 1D modelling simulation performed for SR-Can, see previous
part of Section 3.4.4 and /SKB 2006a, Section 3.4/ shows a good agreement regarding the maximum
permafrost depth at 70,000 years BP (Figure 3-74). The differences observed in Figure 3-74 reflect
the higher thermal conductivity and lower geothermal heat flow values used in /SKB 2006a/ as well
as differences in prescribed surface conditions.
Unfrozen water content and talik formation in reconstruction of last glacial cycle conditions
Figures 3-75 and 3-76 show the unfrozen water content within the frozen bedrock as well as the2D groundwater flow (Darcy velocity) directions at a time of 90.4 kyrs BP for the humid and dryvariants of the reconstruction of last glacial cycle conditions. The figures show the whole profile,and close-ups of the highly conductive deformation zone ZFMA2 at a distance of 4,000 m, and thenear-by area of the two lakes at distances of 9,000 m and 14,600 m. The results show that duringphases of the simulated time period, taliks form at these lake locations (Figures 3-69 and 3-76).
The corresponding results for other time periods are found in /Hartikainen et al. 2010/.
The results show a maximum groundwater flow velocity of ~50 m/yr when no permafrost occurred andsome metres per year beneath the perennially frozen ground. The results show that under continuous
permafrost conditions both the local and regional groundwater flow are reduced considerably in the perennially frozen ground but also in the unfrozen ground beneath permafrost, see Figures 3-75 and 3-76
and /Hartikainen et al. 2010/. On the other hand, the results demonstrate that when the unfrozen groundwater content is greater than 10%, groundwater flow may occur through a continuous partially frozen
permafrost zone, although the flow velocity is very low, only some millimetres per year (Figure 3-76).
This indicate, that under continuous permafrost conditions taliks are able to form under lakes throughperennially frozen ground down to ~50 m depth, if favourable groundwater flow conditions with open
flow paths prevail. When the freezing of ground is advanced to a degree with unfrozen water content
less than 10%, groundwater flow is reduced considerably, and thus a talik is not able to form or survive.
A 2D model approach is in this context sufficient to justify the likelihood of talik occurrence, but not todetermine their groundwater flow and precise geometrical characteristics.

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Salinity concentration in reconstruction of last glacial cycle conditions
Given the permafrost and freezing results presented above, the associated results on salinity concentrations at times of 111.5, 95, 74, 70 kyrs BP are illustrated in Figures 3-77 and 3-78. Due to low salinity
concentrations at shallow depths, the impacts of freezing on salinity exclusion and redistribution are
difficult to see from the results. However, under more severe conditions when the air temperature isdecreased exceptionally, i.e. by 8°C, the freezing occurs more intensively and an increase in salinity
concentration due to the exclusion can be seen /Hartikainen et al. 2010/.
Groundwater flow is an integral part of the salinity transport modelling, which hence suffers fromthe same restrictions as the 2D groundwater flow modelling. However, the results for salinity
concentration (Figure 3-78) and /Hartikainen et al. 2010 Figures 5-32 to 5-35/ show that freezing
can induce salt exclusion and transport when perennially frozen ground development exceeds 200 m
depth. This can be seen as increased salinity concentrations extending downwards in front of thepartially frozen bedrock. At shallow depths the impacts of freezing cannot be seen since the salinity
concentration of the groundwater has been diluted prior to the development of perennially frozen
ground. The increased salinity concentration within the perennially frozen ground is mainly because
salt transport occurs more slowly than the freezing zone advances.
Description and results of sensitivity experiments

Sensitivity experiment 1 – uncertainty in air temperature
In sensitivity experiment 1, the objective is to investigate the impact of the estimated uncertainty
in the air temperature curve. The description of the temperature curve and its estimated uncertainty
is found in Appendix 1. From this description, it is estimated that an uncertainty range of ±6°C isrelevant for the temperature curve used as input. In addition, it was investigated how much the airtemperature curve is required to be lowered to get the 0°C, −2°C and –4°C isotherms to reach the
450 m repository depth. Based on the SR-Can results /SKB 2006a/, temperature shifts of 4, 6, 8, 10,
12, 14 and 16°C were chosen (Figure 3-79). As a pessimistic case, the dry variant of the reconstruction of last glacial cycle conditions was simulated, since it results in deeper permafrost than thehumid variant. The mean thermal properties of the subsurface were used.
0

Depth below ground surface (m)

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Maximum permafrost depth
Humid variant of Repetition of last glacial cycle caseDry variant of Repetition of last glacial cycle case
SR-Can 1D simulation

Maximum depth of perennially frozen ground
350

400
115

Humid variant of Repetition of last glacial cycle caseDry variant of Repetition of last glacial cycle case
SR-Can 1D simulation
110

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95

90

85

Time before present (kyrs)

80

75

70

Figure 3-74. Evolution of permafrost and frozen ground depth over the repository location from the 2D reconstruction of last glacial cycle conditions and from the 1D simulation of “the reference surface conditions of thebase variant” from /SKB 2006a/. The results of the 1D and 2D permafrost models are in good agreement. The
shaded area in blue and red represents the range when considering the dry and humid climate variants of thepresent study. The lilac colour indicates that the results for permafrost and perennially frozen ground overlap.

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90.4 kyrs BP

90.4 kyrs BP, max shown Darcy velocity = 0.44 m/year

90.4 kyrs BP, max shown Darcy velocity = 3.4 m/year

90.4 kyrs BP, max shown Darcy velocity = 26 m/year

Figure 3-75. Unfrozen volumetric (m3/m3) groundwater content contours and Darcy flow directions at
90.4 kyrs BP for the humid variant of the reconstruction of last glacial cycle conditions. The upper panel
shows the whole profile, followed by close-ups of the deformation zone ZFMA2 at ~4,000 m, and the near-by
area of the lakes at the distances 9,000 m and 14,600 m from the south-west side of the profile. Light colour
shows the extent of perennially frozen ground, the red line shows the 0°C isotherm (i.e. permafrost depth),
and the yellow box in the upper panel shows the repository. For further examples of these results, see/Hartikainen et al. 2010/.

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90.4 kyrs BP

90.4 kyrs BP, max shown Darcy velocity = 0.028 m/year

90.4 kyrs BP, max shown Darcy velocity = 0.0037 m/year

90.4 kyrs BP, max shown Darcy velocity = 0.0033 m/year

Figure 3-76. Unfrozen volumetric (m3/m3) groundwater content contours and Darcy flow directions at
90.4 kyrs BP for the dry variant of the reconstruction of last glacial cycle conditions. The upper panel
shows the whole profile, followed by close-ups of the deformation zone ZFMA2 at ~4,000 m, and the near-by
area of the lakes at the distances 9,000 m and 14,600 m from the south-west side of the profile. Light colour
shows the extent of perennially frozen ground, the red line shows the 0°C isotherm (i.e. permafrost depth),
and the yellow box in the upper panel shows the repository. For further examples of these results, see/Hartikainen et al. 2010/.

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111.5 kyrs BP

95 kyrs BP

74 kyrs BP

70 kyrs BP

Figure 3-77. Salinity concentration isolines in (mass-%) at times 111.5, 95, 74, 70 kyrs BP for the humid
variant of the reconstruction of last glacial cycle conditions. The isolines show 0.07, 0.1, 0.2, 0.4, 0.6
and 0.8 mass-% salinity. Blue colour on the top of the profile at 111.5 kyrs BP shows the Baltic Sea
and the light colour the extent of perennially frozen ground. For further examples of similar results, see/Hartikainen et al. 2010/.

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111.5 kyrs BP

95 kyrs BP

74 kyrs BP

70 kyrs BP

Figure 3-78. Salinity concentration isolines in (mass-%) at times 111.5, 95, 74, 70 kyrs BP for the dryvariant of the reconstruction of last glacial cycle conditions. The isolines show 0.07, 0.1, 0.2, 0.4, 0.6, 0.8,
1, 1.2, 1.5, 2 and 3 mass-% salinity. Blue colour on the top of the profile at 111.5 kyrs BP shows the Baltic
Sea and the light colour the extent of perennially frozen ground. For further examples of similar results,
see /Hartikainen et al. 2010/.

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15

10

Change in
air temperature

Temperature (°C)

5

+6 °C
0 °C
−4 °C
−6 °C
−8 °C
−10 °C−12 °C−14 °C−16 °C

0

−5

−10

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−20

−25

−30
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95

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Time before present (kyrs)

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Figure 3-79. Investigated temperature evolutions after shifting the entire reconstructed last glacial cycle
temperature curve by −16, −14, −12, −10, –8, –6, –4, 0 and +6°C compared to the reconstruction of thelast glacial cycle. The 0°C curve constitutes the temperature curve reconstructed for the last glacial cycle.

The evolution of maximum permafrost depth and maximum depth of perennially frozen ground at therepository location considering mean thermal properties and the air temperature changed by –6, 0 and+6°C are shown in Figure 3-80. The evolution of maximum 0°C isotherm depth, i.e. maximum permafrost
depth, and maximum depth of perennially frozen ground over the repository and the whole profile forthe dry variant of the reconstruction of last glacial cycle conditions considering mean thermal properties
and the air temperature changed by –4, –6, –8, −10, −12, −14 and −16°C are shown in Figure 3-81
and 3-82. The evolution of maximum −2°C and –4°C isotherm depth over the repository are shownin Figure 3-83 and 3-84, respectively. It can be seen that temperature shifts of ~8, ~10 and ~14°C are
needed to get the 0°C, −2°C and –4°C isotherms to reach the repository at the time with deepest permafrost at 70 kyrs BP. For details on the setup and results of this experiment, see /Hartikainen et al. 2010/.
The results from the investigation of uncertainties in air temperature show that a 6°C lowering of
the entire temperature curve reconstructed for the last glacial cycle (corresponding to the estimatedmaximum uncertainty in the temperature curve, Appendix 1) increased permafrost depth by almost
60%, from the range of ~180–260 m to ~290–410 m, while a corresponding change of the entire curveof +6°C reduce permafrost depth by 60 to 75%, to a range of ~40–100 m (Figure 3-80). A similar effectcan be seen in the depth of perennially frozen ground, i.e. the temperature curve shift of –6°C results in achange from the range ~190–260 m to ~280–380 m, while a +6°C temperature shift reduces the range to~40–100 m. In addition, the results in Figure 3-81 show that, for the dry variant of the last glacial cycle,
the temperature curve reconstructed for the last glacial cycle needs to be lowered by almost 8°C to makepermafrost reach the repository depth. Furthermore, the results show that a shift of the reconstructed
temperature curve of ~10°C and ~14°C are needed to get the –2°C and –4°C isotherms to reach therepository depth (Figure 3-83 and 3-84). The major part of the difference (of up to more than 100 m)
between results for the repository location and results from the entire profile (Figure 3-81 and 3-82) is dueto the heat from the repository. The effect of topography is minor and further reduced with depth. A lessimportant factor is that some rock domains (RFM026, RFM033, RFM034, RFM040) outside the repository have higher thermal conductivity (2.8%) than the one (RFM029) where the repository is located.

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0Depth below ground surface (m)

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350

Maximum permafrost depth
400
450
500
115

Maximum depth of perennially frozen groundhumid variant
dry variant
humid variant
dry variant
humid variant
dry variant

+6 °C
0 °C
−6 °C
110

humid variant
dry variant
humid variant
dry variant
humid variant
dry variant

+6 °C
0 °C
−6 °C

105

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95

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85

80

Time before present (kyrs)

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70

Figure 3-80. Result of sensitivity experiment on uncertainty in air temperature. Evolution of maximum
permafrost depth (solid lines) and maximum depth of perennially frozen ground (dashed lines) over therepository location considering mean thermal properties and the air temperature curve changed by –6, 0
and +6°C. The shaded areas in blue, green and red represents the range when considering the dry andhumid climate variants. Note that the results for permafrost and perennially frozen ground overlap to alarge degree.

0

Depth below ground surface (m)100Change in
air temperature

200

0°C
−4 ° C

300

− 6°C
− 8°C

400

−10 ° C450

−12 ° C500

−14 ° C
−16 ° C

600

700

800
115

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95

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Time before present (kyrs)

Figure 3-81. Evolution of maximum permafrost depth (black solid line) and maximum depth of perenniallyfrozen ground (black dashed line) at the repository location for the dry variant of the reconstruction of lastglacial cycle conditions. The figure also show corresponding results with the air temperature curve lowered by–4, –6, –8, −10, −12, −14 and −16°C compared to the reconstruction of the last glacial cycle (coloured lines).

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0Depth below ground surface (m)100Change in
air temperature

200

0 °C
−4 °C

300

−6 °C
−8 °C

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−10 °C450−12 °C500−14 °C−16 °C

600

700

800
115

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85

Time before present (kyrs)

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Figure 3-82. Evolution of maximum permafrost depth (black solid lines) and maximum depth of perenniallyfrozen ground (black dashed lines) over the whole profile for the dry variant of the reconstruction of lastglacial cycle conditions. The figure also show corresponding results with the air temperature curve loweredby –4, –6, –8, −10, −12, −14 and −16°C (coloured lines).

0

Depth below ground surface (m)100Change in
air temperature

200

0 °C
−4 °C

300

− 6°C
− 8°C

400

− 10 °C450− 12 °C

500− 14 °C− 16 °C

600
700
800
115

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85

Time before present (kyrs)

80

75

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Figure 3-83. Evolution of maximum −2°C -isotherm depth at the repository location for the dry variantof the reconstruction of last glacial cycle conditions (black line). The figure also show corresponding
results with the air temperature curve lowered by –4, –6, –8, −10, −12, −14 and −16°C compared to
the reconstruction of the last glacial cycle (coloured lines).

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0Depth below ground surface (m)100Change in
air temperature

200

0 °C
−4 °C
−6 °C
−8 °C
−10 °C−12 °C−14 °C−16 °C

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400
450
500
600
700
800
115

110

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Time before present (kyrs)

80

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Figure 3-84. Evolution of maximum –4°C -isotherm depth at the repository location for the dry variantof the reconstruction of last glacial cycle conditions (black line). The figure also show corresponding
results with the air temperature curve lowered by –4, –6, –8, −10, −12, −14 and −16°C compared to
the reconstruction of the last glacial cycle (coloured lines).

Sensitivity experiment 2 – uncertainty in geothermal heat flow
Sensitivity experiment 2 investigates the effects of uncertainties in geothermal heat flow. The meansurface level value of 61 mW/m2 and the range from –14% (minimum) to +12% (maximum) /Sundberg
et al. 2009/ are used for the dry climate variant of the reconstruction of last glacial cycle conditions.
For details on the setup of this experiment, see /Hartikainen et al. 2010/. The results show that a variationof –14% to +12% in geothermal heat flow can cause an approximate fluctuation of +8.0% to –6.3%(~280 m to ~240 m) in permafrost depth, and correspondingly an approximate fluctuation of +7.7% to–5.7% (~260 m to ~230 m) in the depth of perennially frozen ground.
Sensitivity experiment 3 and 4 – uncertainty in bedrock thermal conductivity andthermal diffusivity
Sensitivity experiment 3 was made in order to study the consequences of uncertainty in thermal
conductivity (mean, minimum and maximum). In sensitivity experiment 4, the uncertainty isinvestigated in relation to the mean, minimum and maximum thermal diffusivities. In addition, theeffect of heterogeneity of thermal properties on the development of permafrost and frozen ground isstudied by the mixed thermal diffusivities, which are obtained by using values of highest difference
for adjacent rock domains, see /Hartikainen et al. 2010/. For details on the setup of this experiment,see /Hartikainen et al. 2010/.
The results for thermal conductivity indicated that the mean variation of –6.1% to +8.6% in thermal
conductivity can cause an approximate fluctuation of –4.6% to +2.7% (~250 m to ~270 m) in permafrost depth, and correspondingly an approximate fluctuation of –4.1% to +2.7% (~240 m to ~250 m) inperennially frozen ground depth. The uncertainty in thermal diffusivity causes a slightly larger variationin permafrost depth than solely the uncertainty in thermal conductivity. The results demonstrate thatthe mean variation of –11.5% to +14.6% in thermal diffusivity can cause an approximate fluctuationof –6.0% to +3.4% (~240 m to ~270 m) in permafrost depth, and correspondingly an approximate
fluctuation of –5.5% to +3.4% (~230 m to ~250 m) in perennially frozen ground depth. In addition, thevariation in thermal properties between adjacent rock domains has a minor effect on the development of
permafrost and perennially frozen ground. The reason for this is probably the flat topography of the site,in consequence of which heat transfer mainly takes place vertically. This is also supported by the results
for the ground temperature and perennially frozen ground (Figures 3-68 and 3-69) which show rather
uniform permafrost development with depth when it occurs in continuous form.
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Sensitivity experiment 5 – combination of uncertainties in bedrock thermal properties
In sensitivity experiment 5, the objective was to investigate the combination of uncertainties in dominant
thermal properties of the bedrock. The limits for the uncertainty interval are obtained by combining thethermal properties that are anticipated, based on results from /SKB 2006a/, to enhance the permafrost
development, i.e. maximum thermal conductivity, minimum heat capacity and minimum geothermalheat flow, as well as the thermal properties that are expected to diminish the permafrost development, i.e.
minimum thermal conductivity, maximum heat capacity and maximum geothermal heat flow. The dryvariant of the reconstruction of last glacial cycle conditions was simulated. For details on the setup of
this experiment, see /Hartikainen et al. 2010/.The resulting evolution of maximum permafrost depth andmaximum depth of perennially frozen ground over the repository are shown in Figure 3-85. The results
show that the combined uncertainties can cause a variation of –8.9% to +12.2% (~240 m to ~290 m)
in the permafrost depth and a variation of –6.2% to +14.4% (~230 m to ~280 m) in the depth of
perennially frozen ground.
Sensitivity experiment 6 – combination of uncertainties in surface and bedrock
thermal properties
In sensitivity experiment 6 the uncertainties in subsurface thermal properties from experiment 5 are
combined with the uncertainties in surface conditions (excluding air temperature). For the reconstructionof last glacial cycle conditions, the uncertainty interval is obtained by combining the dry variant withthe thermal properties enhancing permafrost development, and the humid variant of the same casewith thermal properties diminishing permafrost development.
Figure 3-86 shows the evolution of maximum permafrost depth and maximum depth of perenniallyfrozen ground over the repository for the reconstruction of last glacial cycle conditions considering
combined uncertainties in surface conditions and thermal properties as described in /Hartikainen et al.
2010, Section 2.2 and 2.4, and Appendix E, H, I and J/. The calculated uncertainty range for the maximum permafrost depth at 70 kyrs BP is 170–290 m and for the maximum depth of perennially frozen
ground 170–280 m. For details on the setup and results of this experiment, see /Hartikainen et al. 2010/.
0Depth below ground surface (m)

50

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Maximum permafrost depth
300

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115

max. geothermal heat flow + min. thermal diffusivity
min. geothermal heat flow + max. thermal diffusivity

Maximum depth of perennially frozen groundmax. geothermal heat flow + min. thermal diffusivity
min. geothermal heat flow + max. thermal diffusivity
110

105

100

95

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85

Time before present (kyrs)

80

75

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Figure 3-85. Evolution of maximum permafrost depth and maximum depth of perennially frozen ground over
the repository for the dry variant of the reconstruction of last glacial cycle conditions considering combined
uncertainty in bedrock thermal conditions and geothermal heat flow. The shaded area in blue and red represents
the range when considering the dry and humid climate variants. The lilac colour indicates that the results forpermafrost and perennially frozen ground overlap.
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0Depth below ground surface (m)

50

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Maximum permafrost depth
humid variant + max. geothermal heat flow + min. thermal diffusivity

250

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Maximum depth of perennially frozen ground
300

350

400
115

humid variant + max. geothermal heat flow + min. thermal diffusivity
dry variant + min. geothermal heat flow + max. thermal diffusivity

Maximum depth of −2 °C and −4 °C isotherm−2 °C,−4 °C, humid variant + max. geothermal heat flow + min. thermal diffusivity−2 °C,−4 °C, dry variant + min. geothermal heat flow + max. thermal diffusivity

110

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Time before present (kyrs)

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Figure 3-86. Evolution of maximum permafrost depth, maximum depth of perennially frozen ground andmaximum depth of −2 and –4°C isotherms over the repository for the reconstruction of last glacial cycle
conditions considering combined uncertainties in surface conditions and bedrock thermal properties(including geothermal heat flow). The shaded area in blue and red represents the range when considering
the dry and humid climate variants. The lilac colour indicates that the results for permafrost and perenniallyfrozen ground overlap.

Sensitivity experiment 7 – combination of uncertainties in air temperature, surface conditions
and bedrock thermal properties
This case is studied in order to illustrate the effect of setting all known uncertainties to their extremeend (air temperature, surface conditions, geothermal heat flow and all thermal properties) so that theyeither promote or reduce permafrost growth. Note that the case is constructed for illustrational purposes, and that having this combination of parameter settings describes an almost unrealisable case. For
this purpose both the dry and humid variants of the reconstruction of last glacial cycle conditions wereused together with an air temperature change of ±6°C and all thermal properties and geothermal heatflow either enhancing or diminishing permafrost development, see /Hartikainen et al. 2010, Section 2.2
and 2.4, and Appendix E, H, I and J/. Based on the estimates of uncertainty in the air temperature curve(Appendix 1) and on the nature of the conceptual- and data uncertainty associated with the temperature
curve /SKB 2010a/, it is considered adequate to lower the temperature curve reconstructed for the lastglacial cycle (Figure 3-55) by 6° C as a most pessimistic case in the sensitivity analysis for examining
the effect of the temperature curve uncertainty for the reconstruction of last glacial cycle conditions.
Figure 3-87 shows the resulting evolution of maximum permafrost depth and maximum depth of
perennially frozen ground over the repository. The results show, as expected, a very large fluctuation inboth permafrost depth (~40 m to ~460 m) and perennially frozen ground depth (~40 m to ~420 m).
These two cases indicate that if all uncertainties have their most pessimistic settings, which is a quite
unrealistic case, permafrost is able to reach the repository depth of 450 m in the reconstruction oflast glacial cycle conditions.

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Depth below ground surface (m)

100
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115

Maximum permafrost depth+6 °C change + humid variant + max. geothermal heat flow + min. thermal diffusivity−6 °C change + dry variant + min. geothermal heat flow + max. thermal diffusivity

Maximum depth of perennially frozen ground+6 °C change + humid variant + max. geothermal heat flow + min. thermal diffusivity−6 °C change + dry variant + min. geothermal heat flow + max. thermal diffusivity

Maximum depth of −2 °C and −4 °C isotherm−4 °C, −6 °C change + dry variant + min. geothermal heat flow + max. thermal diffusivity−2 °C, −6 °C change + dry variant + min. geothermal heat flow + max. thermal diffusivity
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Figure 3-87. Evolution of maximum permafrost depth, maximum depth of perennially frozen ground andmaximum depth of –4°C isotherm over the repository for the reconstruction of last glacial cycle conditions
considering the unrealistic combination of uncertainties in air temperature, surface conditions, geothermalheat flow and thermal properties all set in their most extreme position favoring permafrost growth.

Sensitivity experiment 8 – repository heat
For illustrational purposes, sensitivity experiment 8 studies the effect of including or excluding theheat generated by the spent fuel. However, the heat from the repository should not be regarded an
uncertainty. Cases with heat and without heat from the repository together with the dry variant of thereconstruction of last glacial cycle conditions were simulated. The resulting evolution of maximum
permafrost depth and maximum depth of perennially frozen ground is illustrated in Figure 3-88. The
results show, as expected, that the heat from the repository reduces the permafrost and perenniallyfrozen ground depths, e.g. permafrost depth is reduced by ~38% at 109,700 years BP and by ~14%
at 70,000 years BP (Figure 3-88). The calculated range for the maximum permafrost depth at 70 kyrs
BP is 259–295 m and for the maximum depth of perennially frozen ground 246–280 m.
Sensitivity experiment 9 – convective heat transfer by groundwater flow
Sensitivity experiment 9 investigates the impact of convective heat transfer due to groundwater flow
and salt transport on the evolution of permafrost and perennially frozen ground. Cases with andwithout 2D groundwater flow and salinity transport in the dry variant of the reconstruction of lastglacial cycle conditions were simulated. The resulting evolution of maximum permafrost depth andmaximum depth of perennially frozen ground over the repository with and without groundwater flow
and salinity transport are shown in Figure 3-89.
The results show that the effect of convective heat transfer due to groundwater flow and salinity
transport on the development of permafrost and perennially frozen ground is very weak. The permafrost depths in consideration of groundwater flow and salinity transport are less than a metre greaterthan those without it (for a permafrost depth of ~260 m). A somewhat larger difference in frozen
depth (~5 metres) is possibly a consequence of reduced salinity concentration due to salt transport.
It should again be emphasised that the results for groundwater flow and salinity are impaired by
the 2D modelling approach. However, the topography of the site is rather flat and does not generatesignificant hydraulic gradients, suggesting that a 3D approach would only have a minor impact on
the results for non-glacial conditions.

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0Depth below ground surface (m)

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Figure 3-88. Evolution of maximum permafrost depth and maximum depth of perennially frozen groundover the repository for the dry variant of the reconstruction of last glacial cycle conditions considering
mean thermal properties and variable heat production from the repository from zero to the reference value,see /Hartikainen et al. 2010/. The shaded area in blue and red represents the range when considering the
dry and humid climate variants. The lilac colour indicates that the results for permafrost and perenniallyfrozen ground overlap.

0

Depth below ground surface (m)

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Maximum depth of perennially frozen groundwithout convection and salt transportwith convection and salt transport
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Time before present (kyrs)

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70

Figure 3-89. Evolution of maximum permafrost depth and maximum depth of perennially frozen groundover the repository for the dry variant of the reconstruction of last glacial cycle conditions considering
mean subsurface thermal conditions, and with and without 2D groundwater flow and salinity transport. The
shaded area in blue and red represents the range within which a result is expected to lie when considering
groundwater flow and salt transport between the reference case and case without groundwater flow and salt
transport. The lilac colour indicates that the results for permafrost and perennially frozen ground overlap.

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A study dedicated at looking at 3D groundwater flow under permafrost conditions at Forsmark, using
a simplified approach to model permafrost, is found in /Vidstrand et al. 2010/. The same study
also includes a comparison between the results from the simplified permafrost solution (as used inDarcyTools), and the 2D permafrost study reported here and in /Hartikainen et al. 2010/.
Main conclusion from the 2D permafrost modelling study
The conclusion drawn from the 1D permafrost investigations /SKB 2006a/ that the surface conditions
can be seen as the driving force for the development of permafrost has been strengthened by the 2D
study. In addition, the conclusion that the subsurface conditions and the heat from the spent fuel act aseither reducing or enhancing factors for the permafrost and perennially frozen ground development isalso supported. The following main conclusions can be made from the 2D permafrost study:•

Given the temperature climate development reconstructed for the last glacial cycle, the study
demonstrates how the site-specific spatial- and temporal development of permafrost and perenniallyfrozen ground takes place at the Forsmark site, developing from sporadic-, to discontinuous- tocontinuous coverage in various combinations and extents.

•

The study provides a full range of sensitivity analyses of uncertainties in sub-surface, and surface
conditions, including the reconstructed input air temperature curve for the last glacial cycle.

•

For the reconstruction of last glacial cycle conditions, the simulated maximum permafrost (the0°C isotherm) depth over the repository is between 180 and 260 m depending on the surface
conditions. The corresponding range for the entire investigated profile is from 220 to 300 m.

•

For the reconstruction of last glacial cycle conditions, the simulated maximum depth of perenniallyfrozen ground over the repository is between 180 and 250 m depending on the surface conditions.
The corresponding range for the entire investigated profile is from 210 to 300 m.

•

The prevailing surface conditions, such as temperatures and surface moisture conditions, arethe main factors governing the spatial and temporal development of permafrost and perenniallyfrozen ground at the Forsmark site.

•

Subsurface conditions, such as bedrock thermal properties, geothermal heat flow and groundwatersalinity, modify the spatial and temporal development of permafrost and perennially frozen
ground, but are of secondary importance for modelled permafrost depths compared with surface
conditions.

•

The variation in thermal properties along the profile as well as 2D groundwater flow only havea slight influence on permafrost (the 0°C isotherm) development. Therefore the uncertainties
introduced by excluding lateral variations in thermal properties, boundary conditions and convective
heat transfer in the 1D model study /SKB 2006a/ are considered insignificant.

•If making the very pessimistic combination of setting all known uncertainties (in air temperature,surface conditions, bedrock thermal conductivity and heat capacity, and geothermal heat flow)in the position most favourable for permafrost growth, permafrost may extend to 450 m depth in
45,000 years in the reconstruction of last glacial cycle conditions. However, it should be notedthat this combination of assumptions is very unrealistic.

•

Under continuous permafrost conditions, the unfrozen groundwater content in the perenniallyfrozen ground under lakes can exceed 10% down to a ~50 m depth. This may indicate that taliks
are able to form under lakes through perennially frozen ground if favourable groundwater flow
conditions with open flow paths prevail. When the unfrozen groundwater content decreases below
10%, groundwater flow is reduced considerably, and taliks are not longer able to form or survive.

•

Freezing can induce salt exclusion and salt transport when perennially frozen ground develops
deeper than ~200 m. At more shallow depths the impacts of freezing are difficult to see since thesalinity of groundwater has been diluted prior to the development of perennially frozen ground.
When salt transport occurs more slowly than the freezing zone advances, the salinity concentration
is increased within the perennially frozen ground.

•

The uncertainty dealing with groundwater flow with this permafrost model remains, since the3D flow network is omitted in the 2D model. However, the topography of the site is rather flat togenerate significant hydraulic gradients, suggesting that this would only have a minor impact on
the results for non-glacial conditions.

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•The model uncertainty regarding the neglect of salinity transport in the 1D model /SKB 2006a/
seems to be insignificant, mainly due to the low rock porosity, although the present 2D model is
not able to describe the groundwater flow realistically.

•

Based on the results and the investigations on ground temperature modelling by /Sundberg et al.
2009/ the uncertainty associated with determination of the ground surface temperature from the
air temperature as well as with estimation of the in situ temperature and geothermal heat flow for
the 1,000–10,000 m-depth for the thermal boundary and initial conditions of the model is reduced
considerably.

Further results from the 2D modelling are found in the severe permafrost case (Section 5.5).

3.4.5 Time perspective
Changes in annual ground surface temperature due to climate change can lead to development of
permafrost and freezing of the ground. If the mean annual ground surface temperature decreases
constantly from the present day value to approximately –8°C, permafrost can develop to the depthof 400 m in a period of 26,000 years at Forsmark (Figure 3-52). Under periglacial conditions, permafrost can aggregate from some centimeters to some decimeters in a year whereas its degradation
can take place several times faster than that (Figure 3-57), especially when the surface temperature isincreased above 0°C and permafrost decays simultaneously from the bottom upwards and from the
top downwards. Large degradation rates can occur also at the onset of glaciation when the surface
temperature rises sharply to the freezing point beneath an advancing ice sheet.
3.4.6 Handling in the safety assessment SR-Site
The last-glacial-cycle evolution of permafrost is investigated by means of numerical modelling. The
results were used to construct parts of the SR-Site reference glacial cycle. A broad range of sensitivity experiments are conducted in order to describe the effects of uncertainties in bedrock, surface
or climate conditions. Two numerical permafrost models were used, a 1D model that was used for
simulations of permafrost at the repository location, and a 2D model that investigated spatial development of permafrost along a profile that crossed the repository site. The permafrost models include
a mathematical expression for freezing and thawing of saline-groundwater-saturated bedrock. The
bedrock is considered as an elastic porous medium and the groundwater as an ideal solution of waterand ionic solvents. The models are based on the principles of continuum mechanics, macroscopic
thermodynamics and the theory of mixtures, and are capable of describing heat transfer, freezing ofsaline water, groundwater flow and deformations of bedrock. To capture the most important factors
and parameters affecting the development of permafrost, sensitivity analyses have been performed
considering the following issues:•••

Surface conditions and climate.
Subsurface conditions.
Presence of the repository.

Surface temperatures, together with the influence of surface covers such as snow, vegetation andwater bodies, have been included as factors of importance in the surface conditions. The investigatedsubsurface conditions are thermal properties of the bedrock and geothermal heat flow. The heat generated by the spent fuel has been included in most simulations. The calculations were carried out intwo steps; 1) by using constant surface temperatures, and subsequently 2) using surface temperatures
based on site-specific climate cases. For further, detailed information on the permafrost modelling,
see the preceding parts of Section 3.4.4.

3.4.7

Handling of uncertainties in SR-Site

Uncertainties in mechanistic understanding

There are no major uncertainties in understanding of mechanistic processes regarding permafrost
development. Minor uncertainties in the 1D modelling experiments are associated with the fact thatthe exclusion of salts in freezing of groundwater is not included. The process of freeze-out of salts is
included in the 2D modelling performed.
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Model simplification uncertainty

The major model simplification in the 1D permafrost modelling study is the exclusion of lateralvariations in physical properties, boundary conditions and geometry. For example, full considerationof the anisotropy of thermal conductivity and heat capacity and the features of water bodies andtopography, as well as the heat generation from the spent fuel, requires 3D modelling. However,this is to a large extent taken care of by the 2D modelling study reported here.
The 1D modelling approach could, in certain situations, result in somewhat higher temperatures thanwould be calculated using a 2D or full 3D approach. In the context of permafrost development, theeffect of groundwater flow, cooling down the bedrock, is the most important factor here. However,compared with heat conduction, groundwater flow has only a minor role in permafrost development, as indicated by the 2D modelling results that included groundwater flow. Furthermore, theanisotropy of thermal properties is not a problem in 1D or 2D, since one can choose a combinationof thermal properties that would give lowest temperatures, or at least very close to the lowest temperatures. Therefore, it is unlikely that 3D simulations would yield notably lower temperatures than
the range obtained by the full series of 1D and 2D sensitivity modelling that have been performed.
The 2D modelling approach used for studying talik formation has limitations since the important 3Dgroundwater flow network is missing.
Input data and data uncertainty

Bedrock data
Some data uncertainty exists when it comes to thermal conductivity and heat capacity of rock at theForsmark site. In the calculation of ground temperature and the rate of freezing, thermal conductivity
is the most important input parameter in terms of thermal properties of the ground. Some uncertainty
also exists in determination of hydraulic and mechanical properties of bedrock and salinity concentrations of groundwater versus depth, see appendices in /Hartikainen et al. 2010/ and references therein.
The in situ ground temperature has been measured in boreholes to a depth between 500 and 1,400 m.
Between different boreholes the temperature deviates in a range of 2°C /SKB 2005a, 2006a, b/.
A considerable uncertainty in the 1D permafrost study is associated with determination of thein situ temperature and geothermal heat flow in the depth range of 1,000–10,000 m for the thermalboundary and initial conditions of the model. Based on the results and the investigations on groundtemperature modelling by /Sundberg et al. 2009/, the uncertainty associated with determination of
the ground surface temperature from the air temperature as well as with estimation of the in situ temperature and geothermal heat flow for the 1,000–10,000 m-depth for the thermal boundary and initialconditions of the model have been reduced considerably in the 2D permafrost modelling study.
The uncertainty in geothermal heat flow was investigated and described in /Sundberg et al. 2009/.
In general the uncertainty in thermal characteristics of the bedrock and geothermal heat flow has asignificantly smaller impact on modelled permafrost and freezing depths than uncertainties related to
ground conditions and climate.
Surface conditions
As a complement to the 1D permafrost studies made for Forsmark (SKB 2006a, Section 3.4 and 4.4.1),
the main objective of the 2D modelling study was to investigate impacts of surface conditions on
the spatial (along the profile) development of permafrost and perennially frozen ground using site
specific information on climate and landscape features including water bodies and topography. An
approach based on TWI was introduced to describe the vegetation and snow cover in relation to theprevailing evolving climate and surface moisture conditions. TWI was used to identify and locate
four relevant surface condition types, dry, moist-humid, wet and peatland, that could be expected
to occur on the profile in Boreal-, Subarctic- and Arctic climate zones. Thereafter, ground surface
temperatures were calculated from an air temperature curve by making use of n-factors that yield astatistical relation between the air and ground surface temperatures in consideration of climate andsurface moisture conditions /Hartikainen et al. 2010/. Due to considerable uncertainties related in
description of surface conditions by TWI and modelling of ground surface temperatures by means of
n-factors, two variants, one humid and one dry for the reconstruction of last glacial cycle conditions
and a severe permafrost case were analyzed.
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Temperature curve
Major uncertainties exist in the temperature curve used for the permafrost modelling, including
uncertainties for its representativity for the last glacial cycle climate as well as for its application todescribe possible future climate cases. A detailed discussion and description of these uncertainties
are found in Appendix 1 and /SKB 2010a/. To cover the estimated uncertainties in air temperature
curve reconstructed for the last glacial cycle, a large range of sensitivity cases and alternative
temperature climate variations of the temperature curve have been analyzed. This is also motivatedby the fact that the variation in surface conditions and climate has a larger impact on modelledpermafrost and freezing depth than bedrock thermal conditions and geothermal heat flow.
Input data for the permafrost modelling are described in /Hartikainen et al. 2010 Appendix H/ and/SKB 2010a/.
In addition, to cover the uncertainty in future climate development, a case with cold non-glaciated conditions, occurring during a significantly longer period than analysed here, is presented in Section 5.5.
Repository heat
In the 1D permafrost study made for SR-Can, repository heat was initiated at 120 kyrs BP, whereas
in the 2D study made for SR-Site it was initiated at 112 kyrs BP. However, this temporal difference
has a negligible effect on the timing of the first permafrost over the repository (see Figure 3-88)
since it is the surface conditions that are the main factor for the initiation and development of permafrost. Furthermore, the difference in timing also has a negligible effect on the maximum permafrost
and freezing depths, since, at the time when permafrost depths are largest (at ~70 kyrs BP), repository heat has declined very much (to ~ 0.3% of the initial value). If heat would have been initiated at
120 kyrs BP in the 2D study, it is estimated that the maximum permafrost and freezing depths would
have been no more than a few metres deeper.

3.4.8 Adequacy of references
The SKB report produced for the handling of permafrost and freezing processes /Hartikainen et al. 2010/
has undergone the SR-Site QA system handling, including a documented factual review procedure.
Also the SR-Can Climate report /SKB 2006a/, from which some of the studies are used, has undergone
QA system handling including a factual review process. Other references used for the handling of
the permafrost and freezing processes, in this report and in /Hartikainen et al. 2010/ , are either peerreviewed papers from the scientific literature or the text books /Yershov 1998, Lide 1999, French 2007/.

3.5

Surface denudation

3.5.1 Overview/general description
The downwearing of the Earth’s surface by exogenic processes is accomplished by weathering, erosion, and transportation of material. The combined effect of all weathering and erosion processes isreferred to as denudation, i.e. denudation is the sum of the processes that results in the wearing awayor the progressive lowering of continental relief. The energy needed for the denudation processes isgained from endogenic and exogenic sources. Glacial erosion is usually regarded as the main erosion
process in Scandinavia during the Quaternary. However, as great parts of Scandinavia were not coveredby ice during long periods of the Quaternary /Porter 1989/, other denudation processes must also
be considered to get the complete picture /Lidmar-Bergström 1997, Lidmar-Bergström et al. 1997,Påsse 2004/.
Weathering

Weathering exerts the most fundamental control on denudation and is the driver of, or limiting
factor, in landscape evolution /Turkington et al. 2005/. Weathering can be defined as structural
and/or mineralogical break down of rock through the cumulative effects of physical, chemical andbiological processes operating at or near the surface, e.g. /Reiche 1950, Selby 1993, Whalley and152

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Warke 2005/. The definition indicates that weathering occurs when minerals/rocks are exposed totemperatures, pressures and moisture conditions characteristic of the atmosphere and hydrosphere,that is in an environment that differs significantly from the conditions in which most igneous andmetamorphic rocks as well as lithified sedimentary rocks were formed. Therefore, the alteration ofrocks by weathering forms new materials (minerals) that are in equilibrium with conditions at ornear the Earth’s surface.
By definition weathering occurs in situ and does not directly involve erosion. This means that it
leads to the formation of a residual material that differs from the parent, unweathered rock withrespect to its physical and chemical properties. Weathering normally lowers the strength of rock andincreases the permeability of the surface material and thus makes it more prone to mass wasting andeasy to erode by running water, glaciers, wind etc. In addition, it is also an important prerequisite forthe widespread development of flora and fauna on land by releasing nutrients for plants and otherorganisms.
Weathering is generally divided into physical-, chemical- and biological components. Physical ormechanical weathering occurs when volumetric expansion and related alteration of stresses lead to
failure and disintegration of the rock. For example, volume changes due to decreased overburden
and stresses can result in the creation of fractures at various scales. Crystallisation and volumetricalteration of salt crystals, freezing of water and freeze-thaw effects, as well as thermal fatigue due torepeated (diurnal) heating and cooling, may also cause physical weathering.
Chemical weathering comprises reactions between rock minerals and water. Examples are solutionof minerals, carbonation, hydrolysis, hydration, and oxidation and reduction. Common to chemical
weathering processes is that they depend on water composition, for example pH, salinity, CO2 andredox potential. The prevailing temperature is another important parameter determining the type andefficiency of chemical weathering.
During temperate conditions in Sweden, weathering only occurs where the bedrock is exposed.
/Swantesson 1992/ has estimated the postglacial weathering of bare bedrock surfaces in southern
Sweden during the Holocene to be less than 0.02 m. Weathering rates in tropical climates range
between 2 and 48 m Ma–1 /Thomas 1994/. Although the weathering rates reported by /Thomas
1994/ have a large span, due to differences in temperature and humidity conditions, it can be seenthat weathering is a slow process even under the most favourable tropical climatic conditions.
Weathering of fresh bedrock under temperate climate conditions is a considerably slower process.
For a description of climatological-, hydrological- and geological factors affecting weathering, see/Olvmo 2010, Section 3.3.3/. Estimated denudation rates for the Forsmark region are presented anddiscussed below.
Erosion

Erosion can be defined as the removal and transport of bedrock and earth materials by a moving natural
agent, such as air, water or ice. Erosion is often preceded by weathering and followed by transport andsedimentation. Air or water flowing over a bed of loose particles generates a shear stress that tends
to initiate particle movement /Collinson 2005a/. There is a critical boundary shear stress related towind velocity and turbulence and water flow velocity above which particle movement occurs. With
the exception of small grain sizes, the critical shear stress will increase as the grain size increases.
For small grain sizes (less than 0.1 mm, silt and clay) increased cohesive strength and lower surface
roughness means that higher velocities are required to initiate movement. In this section, a brief
description is given of the main mechanisms of aeolian-, fluvial- and glacial erosion.
Aeolian erosion
Erosion of sediments by the wind, i.e. aeolian erosion, occurs in environments with sparse or nonexistent vegetation, a supply of fine-grained sediments and strong winds, for example periglacial
regions, semi-arid or arid regions, beaches and agricultural fields. Aeolian processes depend on
weathering or other natural agents, e.g. rivers and waves, to supply sediments for transport. Particle
movement is achieved due to wind shear stress and atmospheric turbulence. Particles begin to movewhen wind forces exceed the effect of weight and cohesion; the first particles to move dislodge or
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impact other grains and the number of particles in movement increase exponentially. However, theparticles reduce the near-bed velocity of the wind and the transport reaches a dynamic equilibrium
state in a few seconds. The mass flux of particles is related to the wind shear velocity, abrasion
and the impact of particles moved by saltation and creep. The latter mobilise fine-grained material
affected by cohesion but easily carried by the wind. Thus, the transport rate is limited both by wind
shear velocity and supply of particles.
Fluvial erosion
Less than 0.005% of the global water is stored in rivers nevertheless they are one of the most, if
not the most, potent erosional forces operating on the Earth’s surface. Rivers cut valleys, transport
sediments and deposit their loads in a variety of depositional environments, such as flood plains anddeltas. Vertical erosion by rivers is a striking feature of the world’s mountainous areas where deep
valleys dissect the landscape and form steep slopes, thereby creating the conditions for mass movement processes that are closely linked to fluvial vertical erosion.
Erosion and deposition of particles by water, i.e. fluvial erosion and deposition, is related to flow
velocity. To generate water flow, a water supply and a flow gradient are required. Therefore, erosion
by water generally requires a slope. However, erosion by water can also occur on bare surfaces asraindrops hit the surface and splash particles away. Fluvial erosion occurs on slopes, in ditches, brooks
and rivers, and on beaches by wave action. Similar to wind erosion, erosion by water depends on shear
stress and turbulence, and particles begin to move when water fluid forces exceed the effects of weightand cohesion. The relation between grain size, flow velocity and different modes of erosion anddeposition are illustrated in a Hjulström-Sundborg diagram, see Figure 3-90.
The material is transported in two distinct ways, it is either carried along with the fluid or in inter­mittent contact with the bed, described in e.g. /Collinson 2005b/. In the first case, fine-grained
particles are supported by the upwards component of turbulence and carried away with the fluid insuspension. In the second case, generally referred to as bedload transport, coarser-grained materialrolls (creep or Erosion reptation) or bounces (saltation) on the bed. Particles carried in suspension
will enhance the erosive capability of the fluid by abrasion and particles moved by saltation, creep
or reptation will set other grains in motion.
Fluvial erosion occurs on slopes provides material that is transported to brooks and rivers, and intolakes and the sea. When freshwater reaches the sea, clay particles can flocculate into larger units andsedimentation is accelerated. Erosion by water on a slope increases with increased intensity of rainor snowmelt, increased inclination and length of the slope and decreased resistance to erosion of thesoil. As for wind erosion, the presence of vegetation reduces the erosion rate.

Suspended load transport

Flow velocity (m/s)

Cr
1

itic
al

er
os

ion

cur
v

D

e
rv
cu
l
na
itio
s
o
ep

Bedload transport

e

0.1

0.01

No transport - deposition
Clay

Silt
0.01

Sand
0.1

Granules1

Grain diameter (mm)Pebbles

Cobbles10

ErosionFigure 3-90. A Hjulström-Sundborg diagram showing the relation between flow velocity, grain size and fluvial erosion.

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Erosion on slopes can also be the result of mud flows, landslides, creep processes and avalanches.
These processes result from the movement of masses of snow, earth and debris down a slope under
the influence of gravity and can occur at scales from a few metres to several kilometres and at rates
from metres per year to metres per second. The material can move by creep, sliding, flowing andfalling. The movement of material is generally triggered by increased water content (for example
from snowmelt), but earthquakes and volcanic eruptions can also cause material to move down a slope.
For more information on fluvial erosion, see /Olvmo 2010, Section 4.1/ and references therein.
Glacial erosion
Glacial erosion is complex and depends on material properties of ice and rocks, ice thermodynamics,
friction and lubrication, chemical effects and subglacial hydrology. In addition, a precondition for
efficient glacial erosion in bedrock is that the bedrock has been subject to pre-glacial sub-aerial
weathering processes.
In recent years, interpretations of geomorphological features e.g. /Kleman et al. 1997, 2008/
measurements of cosmogenic radionuclides, e.g. /Stroeven et al. 2002a/ and ice sheet modelling,e.g. /Näslund et al. 2003/ have resulted in a composite view on glacial erosion over Scandinavian.
In some areas, over time dominated by basal frozen conditions, glacial erosion has been limited,
whereas in other areas, where basal melting and sliding (see Section 3.1) have occurred, erosion may
have been considerable.
Glacial erosion is traditionally divided into two components; plucking and abrasion. Plucking is
the process by which rock fragment of different size are loosened, entrained and transported awayfrom the glacier bed, while abrasion is the grinding of the substrate by rock fragments held in themoving glacier bed cf. /Sugden and John 1976, Benn and Evans 1998/. The former is responsible forthe characteristic irregular and fractured surfaces of lee sides of bedrock bumps and small hills informerly glaciated areas.
Failure leading to loosening of fragments is known to be caused by stresses set up by differential ice loadand high water pressures at the ice-bedrock interface cf. /Sugden and John 1976, Drewry 1986, Benn andEvans 1998/. Many minor fracture features, such as chattermarks and crescentic gouges, observed on
glacially affected rock surfaces, indicate that normal stress at the glacier bed may be sufficient to cause
failure in some rocks. However, the role of pre-existing weakness, such as joints, cracks and foliation, andalso pre-glacial weathering should not be underestimated cf. /Olvmo and Johansson 2002/. In addition,
glacially induced cyclic water pressure variations may be important for bedrock crack propagation e.g.
/Hooke 1991, Iverson 1991/, a process by which rock fragments may be formed and eroded from the
bedrock. Melting of ice on the stoss side and refreezing of meltwater on the lee side of obstacles at the ice
sheet bed may occur under certain thermal and pressure conditions. This so-called regelation process may
also contribute to glacial erosion and transportation of subglacial till e.g. /Iverson 1993, 2000/.
Abrasion is the process whereby the bedrock beneath a glacier is scoured by debris carried in thebasal layers of the glacier cf. /Sugden and John 1976, Drewry 1986/. The process leads to striationand polishing of bedrock surfaces cf. /Benn and Evans 1998/ and is typical of the stoss side of rockbumps in formerly glaciated terrain. Many factors control the effectiveness of glacial abrasion. The
relative hardness of the overriding clast and the bedrock is important and the erosion process is mosteffective when the overriding clast is considerably harder than the substratum. The force pressing theclast against the bed is decisive and depends in turn on the shape of clast as well as the motion of theclast during transport. Other factors related to the availability of clasts in the basal ice layers, such as
clast concentration, removal of debris and availability of basal debris are also of major importance.
Also glaciers and ice sheets that are cold-based could contribute to a restricted amount of glacial
erosion /Cuffey et al. 2000, Waller 2001 Bennett et al. 2003/. This is achieved by e.g. movement,
rotation and entrainment of debris in the basal ice at sub-freezing temperatures. However, thisprocess is of very limited importance compared to other glacial erosional processes.
The magnitude of glacial erosion differs widely both in time and space. At a continental scale,the large-scale pattern of glacial erosion is controlled by ice sheet thermal regime and topographyof the subglacial landscape. Based on a simple glaciological model /Sugden 1977, 1978/ made areconstruction of the thermal regime of the Laurentide ice sheet. The reconstructed thermal pattern

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shows an inner wet-based area and an outer cold-bed area, which broadly corresponds to the patternof glacial erosion indicated by the distribution of erosional landforms. The most intense erosion as
indicated by areas with high lake density coincide with the transition zone between wet-based andcold-based ice in the model, which probably favours plucking and debris entrainment.
/Näslund et al. 2003/ used a numerical ice sheet model to study regional ice flow directions andglacial erosion of the Weichselian ice sheet in Fennoscandia. A new quantity, basal sliding distance,
was introduced, describing the accumulated length of ice that has passed over the landscape by basal
sliding. It was suggested that this entity could be used as a proxy for glacial erosion. The results
indicate high basal sliding distance values in SW Sweden/SE Norway, in Skagerrak, and along the
Gulf of Bothnia, implying relatively large amounts of glacial erosion in these regions. On elevatedparts of the Scandinavian mountain range and on adjacent plains in the east the basal sliding distancevalues are low, implying weaker glacial erosion, which is fairly in agreement with geological andgeomorphological evidence cf. /Lagerbäck and Robertsson 1988, Riis 1996, Stroeven et al. 2002b,
Olvmo et al. 2005/. The method of estimating glacial erosion by simulated basal sliding distance/Näslund et al. 2003/ was further developed by /Staiger et al. 2005/ who introduced a normalizationof the sliding values by the duration of ice cover over a particular site. /Staiger et al. 2005/ also set
up a relationship between normalized sliding distance and rate of glacial erosion.
Another approach to the issue of glacial erosion is presented by /Hallet et al. 1996/ who made a comprehensive review of glacial erosion rates based on sediment yields. They found that rates of glacial
erosion vary by many orders of magnitude from 0.01 mm yr−1 for polar glaciers and thin temperate
plateau glaciers on crystalline bedrock, to 0.1 mm yr−1 for temperate valley glaciers also on resistant
crystalline bedrock in Norway, to 1.0 mm yr−1 for small temperate glaciers on diverse bedrock in the
Swiss Alps, and to 10–100 mm yr−1 for large and fast-moving temperate valley glaciers in the tectonically active ranges of southeast Alaska. These major differences highlight the importance of the glacial
basal thermal regime, glacial dynamics and topographic relief on the rates of glacial erosion.
Yet another approach was presented by /Påsse 2004/. In order to estimate the average glacial erosionin the bedrock in non-mountainous regions he used seismic data and well depth data on the thicknessof the minerogenic Quaternary sediments in Sweden and Denmark. The average thickness of
Quaternary sediments was estimated to be 16 m in the investigated area, which corresponds to 12 mof bedrock assuming that the whole volume is the result of glacial erosion of fresh bedrock. Sincea great part of the sediments likely consist of glacially redistributed Tertiary regolith, this number
probably is an overestimation of the glacial erosion depth in the bedrock. Considering this, Påsse
concluded that the average glacial erosion during a full glacial period may be estimated to between
0.2 m and 4 m. This is in agreement with estimates of glacial erosion in the Precambrian basement
based on geomorphological observations /Lidmar-Bergström 1997, Ebert 2009/. Lidmar-Bergström
distinguishes the estimates of glacial erosion of Tertiary saprolites from glacial erosion of fresh bedrock. The glacial erosion of saprolites is estimated between 10 and 50 m and glacial erosion of fresh
bedrock is estimated at some tens of metre /Lidmar-Bergström 1997/, although with great variations.
However, in Fennoscandia as a whole, large spatial differences in thicknesses of Quaternary deposits
occur and distinct patterns of glacial scouring and deep linear erosion are observed in places.
/Kleman et al. 2008/ point to the relative roles of mountain ice sheets and full-sized Fennoscandian ice
sheets for this zonation and use spatio-temporal qualitative modelling of ice sheet extent and migrationof erosion and deposition zones through the entire Quaternary to suggest an explanatory model forthe current spatial pattern of Quaternary deposits and erosion zones.
Glacial meltwater erosion may be an effective agent both in subglacial and proglacial environments.
The sediment concentrations of glacial meltwater streams are often high and the flow is often very
rapid and turbulent, which mean that flows transitional between debris flows and normal stream flow
are common /Benn and Evans 1998/. The erosivity of glacial streams is therefore often high both on
bedrock and sediments. Apart from the high erosivity, the mechanisms of glacial meltwater erosion
are the same as normal fluvial erosion including abrasion, cavitation, fluid stressing and particle
entrainment from cohesionless beds as well as chemical erosion.
The relative efficacy of fluvial and glacial erosion, on a global basis, are presented in /Koppes andMontgomery 2009/. For a more detailed description of erosional and weathering processes relevant
for Scandinavian conditions and the Forsmark area in a 0.1 to 1 Myr time perspective, see /Olvmo
2010/. Estimated denudation rates for the Forsmark region are presented and discussed below.
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3.5.2 Controlling conditions and factors
As mentioned above, weathering and erosion are complex processes affected by the prevailingenvironmental conditions. Given the characteristics of the rock, the degree of weathering depends on
the availability of weathering agents such as salt, moisture, biota (e.g. microorganisms and lichens),and on the microclimatic conditions to which the rock is exposed. Weathering occurs in all climates,
but high temperatures and humid conditions increase weathering rates. In areas where frost and/or
permafrost occur, freezing and freeze-thaw effects are important geomorphic processes. However,weathering in periglacial environments is also affected by salt (in combination with freezing), wetting
and drying, thermal fatigue and biological agents /French 2007/.
During temperate climate conditions, the dominant process is fluvial erosion. During permafrost
climate conditions, when the climate can be expected to be dryer and vegetation sparse, aeolian
erosion may also be important. During spring when the snow melts, there may be significant erosionof sediments in the active layer in periglacial permafrost areas. During glacial climate conditions, theerosion is mainly affected by the basal conditions of the ice sheet. In areas where the ice is frozen
to the bed, the ice acts to preserve its subsurface and practically no erosion occurs, whereas in areasof basal melting erosion of bedrock and pre-existing sediment is likely to occur, with more erosion
generally associated with faster ice flow. In sub-glacial tunnels, where melt water flows towards theice sheet margin, erosion of loose sediments and bedrock can be significant.
The topographic relief in Scandinavia is generally not dramatic outside the Scandinavian mountains,e.g. /Lidmar-Bergström and Näslund 2005/. The resistance to denudation of the Precambrian rocks
that dominate in Scandinavia is rather uniform, as compared with, for example, the variation between
the Caledonian bedrock types in the Scandinavian mountains. Although the relief locally may be
steep at many places, the relatively homogeneous denudation resistance has contributed to the moderate to low relief of the Precambrian rock basement. The relative relief is exceptionally low (less thansome tens of metre) along significant portions of the eastern Baltic Sea coastline of southern andsouth-central Sweden, due to the presence of the sub-Cambrian peneplain (a denudation surface
formed in crystalline rock with a typical relief of less than 20 m /Rudberg 1954, Lidmar-Bergström
1995/, interpreted to have been formed down to sea-level prior to the Cambrian). The low relief
favours low subaerial erosion rates, for example, the most important subaerial erosion process,fluvial erosion, increases with steeper topographical gradients.
One main reason why a deep dissection of the landscape could not occur by fluvial erosion in Forsmark
is because of the coastal location of the site, situated just a few m above the Baltic Sea. The sea-levelacts as the base-level for fluvial erosion, i.e. fluvial erosion cannot cut deeper than sea-level. The
expected maximum lowering of the Baltic sea-level during glacial cycles (18 m, see Section 4.5.2) sets,
together with the present altitude of the ground surface above the repository, a maximum limit to theamount future fluvial erosion at the Forsmark site.
Table 3-16 summarises how geosphere variables are influenced by surface denudation.

3.5.3 Natural analogues/observations in nature
The current landforms in Scandinavia are a result of the interaction between tectonic processes, weathering, erosion and sedimentation and consequently the traces of these processes can be observed andinterpreted in nature. In Scandinavia, large-scale bedrock landforms outside the mountain range, such as
denudation surfaces and landscapes of weathering and stripping of saprolites, are common, and reflect
the varying phases of weathering, erosion, and sedimentation that have occurred in association withtectonic events and climate change e.g. /Lidmar-Bergström 1997, Lidmar-Bergström and Näslund 2002/.
During the late Cenozoic glacial cycles, ice sheets have repeatedly covered parts or the whole of
Scandinavia, producing glacial erosional landforms in bedrock e.g. /Rudberg 1954/ and in surficial
deposits e.g. /Kleman et al. 1997, 2008/. Traditionally, it has been thought that a considerable part ofthe bedrock relief is the result of glacial erosion. However, as mentioned above, it has been shown thatthe amount of glacial erosion of the Precambrian basement in general is on the order of a few tens ofmetres /Lidmar-Bergström 1997/, and in this context valleys are more eroded than the surrounding terrain. However, within the Scandinavian mountain range, valleys may have experienced several hundred
metres of glacial erosion /Kleman and Stroeven 1997/. /Lidmar-Bergström 1997/ further estimatedTR-10-49

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Table 3-16. Influence of surface denudation on geosphere variables.
Geosphere variable

Influence present (Y/N)

Summary of influence

Groundwater flowNo

The process refers to weathering and erosion at thesurface. Even if there is no absolute boundary between
surface- and groundwater this process only includes
the action of surface water.

Groundwater composition

Yes

Surface weathering may to some degree change
groundwater composition. The effect is judged as
negligible for the safety assessment.

Rock stresses

Yes

The removal of bedrock by denudation changes thestress field. However, since the denudation rate is very
low, the influence is judged as negligible for the safetyassessment period.

Fracture geometry

Yes

Weathering may change fracture aperture in theupper metre of bedrock, for instance by widening of
fractures due to freezing. The influence is judged as
negligible in relation to repository safety.

that denudation occurring during the Proterozoic (600–2,500 million years ago) was on the order oftens of kilometres, and during the Mesozoic (65–245 million years ago) and Tertiary (2.5–65 million
years ago) up to 600 m at the most. Furthermore, during the last decades our understanding of ice
sheet thermodynamics has showed that large portions of ice sheets may be cold-based and thus have anegligible erosive effect. Over the same period, numerous glacial geological studies have shown thatthe occurrence of preglacial saprolite remnants is solid evidence that ice sheets under certain circumstances cause very restricted glacial erosion, or even have a preserving effect on pre-glacial landforms/­Lagerbäck 1988a, b/ and /Kleman 1994/, whereas under other conditions they may be fully erosive.

3.5.4 Model studies
Estimates of denudation rates can be made either through numerical models that per se describe
the processes involved in the denudation, for instance the processes resulting in glacial erosion, or
estimates may be based on interpretation and GIS-modelling of the development over time of landforms. A review of numerical modelling of glacial erosion is found in /Fisher 2009/, while the latterapproach may be exemplified by /Olvmo 2010/, who has used this method in a study of long-term
denudation history of southern Sweden.
The long-term denudation history of southern Sweden is interesting for two reasons. Firstly, it may be
a guide to understand the denudation rate through time in the Forsmark area. Secondly, it is importantbecause the different landforms and surfaces that have been recognized, such as the sub-Cambrian
peneplain mentioned above, may be used as reference surfaces to understand, at least at a regional
scale, the magnitude and patterns of glacial erosion, which is a important issue with respect to the time
perspective of the safety assessment.
Some examples from Sweden of calculated long-term denudation rates can be found. Many of them
use remnants of the sub-Cambrian peneplain as a reference surface. In eastern Småland and southernÖstergötland, the maximum denudation of the tectonically uplifted sub-Cambrian peneplain is approximately 100 m. Once overlying sedimentary cover rocks had been removed, denudation and removal ofthe denudation products would have required a period of between 2 and 50 Myrs /Lidmar-Bergström
et al. 1997/. On the Swedish west coast, a so-called Sub-Cretaceous etch surface has been incised into
the sub-Cambrian peneplain by weathering. In places, it has resulted in a maximum bedrock loweringof 135 m. The time required for this amount of denudation has been suggested to be between 3 and
68 Myrs /Lidmar-Bergström et al. 1997/, with the denudation occurring during the Jurassic-Early
Cretaceous. Finally, saprolite thicknesses of 50 m in Skåne were interpreted to have been produced byweathering over a time interval of 1–25 Myrs /Lidmar-Bergström et al. 1997/.
Given the specific tectonic and climatic evolution, including phases of denudation, sedimentation,and stripping of saprolites, the above examples clearly show that the evolution of the sub-Cambrian
peneplain in southern Sweden into a younger incised landscape has been a very slow process, and thatthe total amount of material removed from the crystalline bedrock has not been large considering the158

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very long time periods involved. The total maximum denudation of bedrock in areas of Precambrian
crystalline basement in southern Sweden during the last 540 Myrs is on the order of a few hundred
metres. A significant part of this denudation took place during Mesozoic tropical climate conditions,considerably more favourable for chemical deep weathering than for instance a temperate climate.
A more detailed account of the long-term landform development and associated denudation insouthern Sweden is given in /Olvmo 2010/.
A study on landform development and associated denudation in the Forsmark and Laxemar regions
has been performed, reported in /Olvmo 2010/. In the following, a summary of the study in the
Forsmark region is given.
From a geomorphological point of view the investigated region (Figure 3-91) can be divided into four
distinctive areas, i) a central area, hosting the Forsmark site, characterized by very low topographic
relief, ii) a coastal region SE of Forsmark with coast-parallel lineaments, iii) a southern part with more
pronounced relief characterized by east-west trending fault scarps, and iv) an area with hilly relief
in the northwest. The relief in the central area, including the Forsmark site, is characteristic of the
sub-Cambrian peneplain and is consistent with the interpretation by /Lidmar-Bergström 1997/. The
relative relief in this area is less than 20 m (Figure 3-91), although with some shallow valleys present
in the peneplain. Around Forsmark (within c. 10 km) the relief is extremely low, often less than 10 m.
South of this central flat area, a more dissected landscape appears. Here the peneplain is broken into
blocks, tilted in different directions. Some of the blocks are elevated and partly dissected by weatheringand erosion. In the southern part of the study area, the elevated rims of these uplifted blocks give rise toeast-west trending horst ridges, whereas in the Stockholm region the peneplain is highly dissected andgive rise to a joint valley landscape, consisting of plateaux bounded by shallow straight, structurally
controlled valleys. In the coastal low lying areas between Forsmark and Uppsala, the relief is quiteconsiderable along coast parallel tectonic lineaments. This is obvious in the relative relief map(Figure 3-91) where relative relief locally rises to 50 m along these fracture zones. In this area,some 15 km southeast of Forsmark, the coastal landscape is more dissected. The area coincide withareas with a high frequency of rock outcrops which may suggest that glacial erosion has been more
effective in this region than at Forsmark, probably as a result of the closeness to the Baltic Sea
depression which may have influenced the ice sheet flow.
The Forsmark area is situated on the sub-Cambrian Peneplain not far from the present extension of
the Lower Palaeozoic sedimentary cover rocks. Exhumation of the peneplain has probably occurred
during the Pleistocene glaciations and possibly as late as the Weichselian glaciation as indicated by
the occurrence of clayey tills in the area. The major bedrock landforms in the area are probably close
to the original peneplain surface. The area is extremely flat over large distances and many landforms
in the Quaternary drift cover are prominent features in the landscape.
In this context it is worth noting that the relief map (Figure 3-91) shows the ground surface relief, i.e.
it comprises the relief of both bedrock and overlying Quaternary deposits. The uncertainty introduced
by not having removed these deposits in the analysis is however in general not large, exemplified by
the fact that within the terrestrial parts of the Forsmark site (Figure 3-59), the average soil thickness
is only 4 m /SKB 2010e Section 5.1.2/. Locally in the investigated area (Figure 3-91), larger sediment
thicknesses of up to c. 100 m may occur, typically where large eskers are found. However, parts ofthese sediment thicknesses do have an expression in the relief map, such as the north-south trending
eskers located north of Uppsala (Figure 3-91). If the esker sediments could be removed from the DEM,
the resulting relative relief would not necessarily be greater than the one of the present eskers.
Nevertheless, this uncertainty is taken care of by the analysis of a case with considerably higher
relief than in Figure 3-91, see below.
Glacial erosion has reshaped some of the forms in the bedrock and may be responsible for evacuationof regolith along structurally controlled valleys. This is especially true in the coastal areas some 15 km
southeast of Forsmark, where glacial erosion may have been comparably effective along fracture zones.
Based on the conclusion that the sub-Cambrian peneplain is well-preserved and probably coincident
with the present relief the total amount of glacial erosion caused by repeated Pleistocene glaciationsin the area on average is probably less than 10 metres, but is probably above that number in the coastalzone southeast of Forsmark. The very low relief of the landscape today, together with the fact that the
bedrock consists of crystalline basement rocks, implies that the present denudation rate is extremely
low and corresponds to the low rates found in low relief shield areas reported in the literature.

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Figure 3-91. Relative relief of the ground surface in the Forsmark study area.

Future glacial erosion at Forsmark

The effect of a future glacial cycle would probably be very limited in areas dominated by the wellpreserved sub-Cambrian peneplain, such as at Forsmark. Given the long-term denudation history
up to present, the expected amount of glacial erosion during a future glacial cycle is probably very
limited. For the Forsmark region, i.e. an area significantly larger than the specific repository location, /Olvmo 2010/ estimated that glacial erosion could amount to 2–5 m for a future glacial cycle
comparable with the last one. /Lidmar-Bergström 1997/ estimated the total amount of glacial erosion
for all Late Cenozoic ice sheets in low-land terrain in Sweden to be around 10 m. If extreme assumptions are excluded, which is valid approach for the low relief Forsmark site, about 1 m of glacial
erosion for one glacial cycle was estimated by /Påsse 2004/. Based on all these estimates, and the
topographic and geological characteristics of the repository location, the amount of glacial erosion in
fresh bedrock for one glacial cycle is here estimated to 1–2 m at the repository location. More glacial
erosion is expected in topographic low positions and/or along major fracture zones, especially ifsuch features are located along the general ice flow direction. The dissected coastal areas south-eastof Forsmark, with sometimes considerable relief, may experience stronger glacial erosion during
a glaciation with a similar flow pattern as during the Late Weichselian. In this area, c. 15 km from
Forsmark, glacial erosion of more than 10 m in one glacial cycle would be expected locally in low
topographic positions /Olvmo 2010/.
Future non-glacial contribution to denudation at Forsmark in a 1 Myr time perspective

In order to evaluate the potential denudation in the Forsmark region in a 1 Myr time perspective, apotential denudation rate is calculated based on the relief of the area. The calculation is based on an
empirical relationship between denudation and relief /Ahnert 1970/. The calculation excludes theeffect of land uplift and glacial erosion and simply relates denudation to the local relief assuming
D= 0.1535h								Equation 3-12
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where D is the denudation rate, h is the relative relief and is a substitute for mean slope. The map ofrelative relief (Figure 3-91) was used for the calculation. The results show that the potential future,
non-glacial component of denudation is between 0–5 m/Myr over most of the area (including theForsmark site) (Figure 3-92), but might rise to between 5 and 10 m/Myr along uplifted fault scarps
in the south, and to 10–25 m/Myr in the northwest where the contemporary relief is higher. The
denudation rates resulting from this type of calculation are verified by a comparison of corresponding
denudation rates for the Laxemar region with denudation and relief formation induced by the upliftof the so-called South Swedish Dome in the late Tertiary /Olvmo 2010/.
In order to show the effect of a change in relative relief by, for instance, a tectonic uplift occurring
many millions of years into the future, a map of potential denudation after a five-fold increase of therelative relief was produced. This case also covers the uncertainty introduced by not having excluded
sediment thicknesses from the relief map. This case does not significantly change the picture for theForsmark site. Even if the local relief were raised to above 100 m, for instance by tectonic movement,
the estimated non-glacial denudation is very low, up to 5–20 m/Myr /Olvmo 2010/.
Conclusions

The Forsmark area has a very low topographic relief and the repository location is not situated ina major fracture zone, resulting in a setting not prone to effective denudation. Given this, and thedenudation rates estimated by /Lidmar-Bergström 1997, Påsse 2004, Olvmo 2010/, a glacial erosion
value of 1–2 m per glacial cycle (similar to the last one), is used in the summary below. If looking
only at the non-glacial component of denudation, the long-term denudation rate of the site is low as aconsequence of the very low relief and the proximity to base level (sea-level). The numbers estimatedfor the long-term denudation rates /Olvmo 2010/ are in agreement with reports of denudation rates ingeological shield areas and lie within the range 0 to 10 m/Myr.
A summary of expected future denudations rates at the Forsmark repository location is provided inTable 3-17. Locally, in topographic depressions away from the repository location, such as alonglarger valleys and fracture zones, denudation values may be higher than this. However, the Forsmark
repository site is not located in a topographic low, nor in a major fracture zone, and therefore theestimated values in Table 3-17 are regarded as valid for the repository location.

3.5.5 Time perspective
Weathering, erosion and sedimentation are active processes over most of the Earth’s surface.
In Sweden, the maximum total denudation can be expected to be limited in lowland areas of low
relative relief. For the region around Forsmark it is estimated to be up to 20 m Myr–1. The rate ofweathering and erosion vary in time and space.
3.5.6 Handling in the safety assessment SR-Site
Present-day knowledge on denudation processes suggests that denudation generally will be limited
in lowland parts of Scandinavia during the assessment period. Areas where the bedrock is highly
fractured and weathering and erosion can be expected to be significantly more rapid than the estimatedaverages do not apply for the repository location at Forsmark. The impact of weathering and erosion on
groundwater flow and rock stresses is considered to be insignificant over the time period of relevancein assessments. In a very long time perspective, several million years, weathering and erosion may
alter topography and rock stresses by redistribution of sediments and rock mass. In an even longer time
perspective, many tens or hundreds of millions of years or longer, the repository may be brought to theground surface due to tectonic uplift in conjunction with a combination of denudation processes.
Current knowledge, e.g. /Thomas 1994, Lidmar-Bergström 1997, 2007, Kleman 1994, Påsse 2004,
Olvmo 2010/, supports the conclusion that the amount of future surface denudation is insignificant
for repository safety in the time perspective of 100,000 years and one million years. The amount ofdenudation is estimated to be ~20 m or less for the Forsmark repository location during the coming
1 Myr (Table 3-17).

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Figure 3-92. Calculated potential future non-glacial denudation in the Forsmark region. At Forsmark,calculated values of the non-glacial denudation component are low, up to 5 m/Myr. From /Olvmo 2010/.
Table 3-17. Expected future denudation at the repository location in Forsmark in a 100 kyr and
1 Myr time perspective. If looking at a larger region around Forsmark, locally higher denudation
values are expected in topographic low positions and along major fracture zones. However, theForsmark repository location is not located at a topographic low, nor in a major fracture zone,
and therefore the estimated values are regarded as valid for the repository.
Parameter~100 kyr timescale(one glacial cycle)

1 million year timescale(~ 8 glacial cycles)

A. Glacial erosion(average value over landscape)

1–2 m

8–16 mB. Non-glacial denudation(0–0.6)0–5 m

C. Total denudation (A+B)

1 to 2.6 m

8–21 m162

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3.5.7 Handling of uncertainties in SR-Site
Rates of weathering and erosion vary as environmental conditions change. The rate of weathering anderosion also depends on rock and sediment characteristics and topography. In spite of the uncertainties
in future environmental conditions, there is no reason to believe that denudation in the Forsmark
region would diverge significantly from the values estimated above. The area has a low relief with
low topographical gradients, in turn producing limited fluvial and glacial erosion at the site.
Uncertainties in mechanistic understanding

There are uncertainties in the understanding of the detailed mechanisms of weathering and erosion.
However, for the safety assessment, the approach of using also indirect estimates of the amount oferosion and weathering over long time-scales (i.e. the development of large-scale landforms), thetotal knowledge is sufficient for estimating the amount of weathering and erosion of fresh bedrock
during the assessment period.
Model simplification uncertainties in SR-Site

Not relevant for this process.
Input data and data uncertainties in SR-Site

Not relevant for this process.

3.5.8 Adequacy of references
The SKB report produced for the handling of surface denudation /Olvmo 2010/ has undergone the
SR-Site QA system handling, including a documented factual review procedure. Also the SR-Can
Geosphere process report /SKB 2006b/, from which some of the studies are used, have undergone
QA system handling including a factual review process. Other references used for the handling of
surface denudation processes are either peer-reviewed papers from the scientific literature or fromtext books /Sugden and John 1976, Drewry 1986, Paterson 1994, Thomas 1994, Benn and Evans
1998, French 2007/.

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