The interglacial–glacial cycle and geochemical evolution of Canadian and Fennoscandian Shield groundwaters

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Geochimica et Cosmochimica Acta 76 (2012) 45–67 www.elsevier.com/locate/gca

The interglacial–glacial cycle and geochemical evolution of Canadian and Fennoscandian Shield groundwaters R.L. Stotler a,⇑, S.K. Frape b, T. Ruskeeniemi c, P. Pitka¨nen d, D.W. Blowes b b

a Department of Geology, University of Kansas, 1475 Jayhawk Blvd., Lawrence, KS 66045, USA Department of Earth Sciences, University of Waterloo, 200 University Ave. West, Waterloo, ON, Canada N2L 3G1 c Geological Survey of Finland, Espoo, Finland d Posiva Oy, Eurajoki, Finland

Received 30 November 2010; accepted in revised form 3 October 2011; available online 10 October 2011

Abstract Results from cryogenic column experiments are compared with the geochemical data collected in the Canadian and Fennoscandian Shields over the past 25 years to investigate the relative influence of the glacial–interglacial cycle; specifically, the impact of continental glaciers, permafrost, and methane hydrate, on the evolution of groundwater from crystalline shield environments. Several different geochemical indicators of freezing processes (either glacial or permafrost-related) were utilized: comparisons of Na/Cl and Br/Cl ratios, d18O and d2H values, and d18O values and Cl concentration. During freezing, fluids with different dominant cations follow distinctly different linear trends when Na/Cl and Br/Cl ratios are compared. Significantly, none of the freezing trends follows the trend hypothesized by Herut et al. (1990) for the evolution of seawater chemistry during freezing. Intrusion of glacial meltwater and in situ freezing (i.e., permafrost formation) result in a similar end-member when comparing d18O values and Cl concentration. The geochemical influence of a freezing process on fresh, brackish, and some saline fluids was identified at some, but not all Canadian Shield sites, regardless of site location with respect to modern-day permafrost. Appreciably, physical and geochemical data do not support the formation of brines through any freezing process in the Canadian and Fennoscandian Shields, as hypothesized by Starinsky and Katz (2003). Rather, on all diagnostic freezing plots, brines are an end-member, indicating a different evolutionary pathway. Significant depletions in 18O with respect to modern precipitation, an indication of either glacial meltwater or a freezing process, were identified at depths of up to 1 km at some sites in the Canadian Shield, and to shallower depths in the Fennoscandian Shield. The potential of this fluid to reach such depths could be attributable to artificial gradients and mixing, glacial recharge, permafrost or paleo-permafrost formation, or methane hydrate or paleo-methane hydrate formation. At most locations it was not possible to distinguish between the different scenarios using the current geochemical database. Ó 2011 Elsevier Ltd. All rights reserved.

1. INTRODUCTION Cold climate processes have affected the evolution of deep fluids far beyond arctic regions. Glacial recharge to groundwater was first described after a region of dilute

⇑ Corresponding author. Address: Department of Geology, Lindley Hall Rm. 120, University of Kansas, 1475 Jayhawk Blvd., Lawrence, KS 66045, USA. Tel.: +1 785 864 6048; fax: +1 785 864 5276. E-mail address: [email protected] (R.L. Stotler).

0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2011.10.006

water with anomalous d2H and d18O isotopic values was observed in aquifers located in the central United States (Perry et al., 1982; Siegel and Mandel, 1984). Since then, regional scale flow reversal, hydrodynamic blowout structures and dissolution of evaporates, biodegradation of oil and bacteriogenic methane formation, dilution of formation waters, and crystalline shield brine formation have all been attributed to glacial influence on subsurface fluids (Siegel, 1991; Bein and Arad, 1992; Martini et al., 1998; Clark et al., 2000; Douglas et al., 2000; Grasby et al., 2000; McIntosh et al., 2002; Starinsky and Katz, 2003; McIntosh and Walter, 2005, 2006; Person et al., 2007).

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Continental ice sheets comparable to the Wisconsinan in North America occupied Scandinavia and the northern part of Central Europe during the Weichselian Ice Age. Extensive intrusion of glacial meltwater into proglacial sedimentary aquifers and the formation of meltwater flow channels have been described, e.g. from Estonia (Vaikma¨e et al., 2001; Raidla et al., 2009) and Germany and Poland (Pietrowski, 1997, 2006). Less pronounced, but still clear evidence of meltwater intrusion into several hundreds of meters of fractured Fennoscandian Shield rocks has been reported in several locations in Finland and Sweden (Smellie and Frape, 1997; Glynn et al., 1999; Laaksoharju and Rehn, 1999; Blomqvist et al., 2000; Pitka¨nen et al., 2001, 2004). During Plio-Pleistocene glaciations, permafrost often formed in front of the advancing ice sheets. At the last glacial maximum, continental glaciers advanced well beyond the extent of the Canadian and Fennoscandian Shields outcrop, with permafrost extending even further (Figs. 1 and 2, Dyke and Prest, 1987; Dyke et al., 2002; Tarasov and Peltier, 2007). Modern permafrost depths in previously unglaciated portions of Siberia are >1580 m, but extend to <700 m in previously glaciated terrain in North America (Natural Resources Canada, 1995; Alexeev and Alexeeva, 2002, 2003). Glacial climate models indicate the maximum permafrost depth over the last 120 ky across North America were similar to present-day values, with repeated glacial advances melting some, if not all, of the permafrost (Tarasov and Peltier, 2007; Lemieux et al., 2008a,b,c). Along the southern extent of the Canadian Shield, where

Fig. 2. Ensemble mean permafrost depth difference between the last glacial maximum (20 kya) and present. Modified from Tarasov and Peltier (2007).

permafrost is currently absent, permafrost depths probably did not exceed 40–60 m during this period (Fig. 2, Tarasov and Peltier, 2007). Thus during Plio-Pleistocene glaciations,

Fig. 1. Map of Canadian Shield and Fennoscandian Shield sample sites. Canadian Shield sites are shown in relation to permafrost distribution and maximum glacial extent at the last glacial maximum (dotted line). Permafrost distribution map modified from Natural Resources Canada (1995). Glacial extent from Dyke et al. (2002). At the last glacial maximum, Finland and Sweden were both completely glaciated.

Freezing processes and evolution of Shield groundwater

permafrost was much shallower or absent in areas currently affected by permafrost, but was present at latitudes much further south than at present (Fig. 2). The potential for permafrost to affect the evolution of deep continental groundwater and/or gas chemistry has only recently been considered. Very few direct investigations of groundwater beneath deep permafrost (>300 m) have been undertaken; most were focused on the sedimentary units of the Siberian Platform (e.g. Pinnekar, 1973; Alexeev and Alexeeva, 2002, 2003; Shouakar-Stash et al., 2007). There, increased concentrations of carbonate, sulfate, sodium, magnesium, and calcium in subpermafrost groundwater were attributed in part to the “freeze-out” process, although other processes, including evaporation and halite dissolution, also influenced these fluids (Alexeev and Alexeeva, 2002, 2003; Shouakar-Stash et al., 2007). A series of gas hydrate research boreholes drilled in the sedimentary sequences of the Mackenzie Delta, Canada, has also provided information about deep fluids in permafrost-affected regions (Dallimore et al., 1999; Dallimore and Collett, 2005). Gas hydrates or clathrates, form in certain low temperature and high pressure conditions, in areas with high methane concentrations, including permafrost and deep sea environments (Kvenvolden and Lorenson, 2001). Recently, groundwater and gases were characterized at the Lupin Mine, Canada, within a part of the Canadian Shield affected by 540 m of permafrost (Ruskeeniemi et al., 2002, 2004; Stotler et al., 2003, 2006, 2008, 2009, 2010b; Lambie et al., 2004; Ne´grel et al., 2004; Casanova et al., 2005; Ne´grel and Casanova, 2005; Ahonen et al., 2008; Stotler, 2008; Onstott et al., 2009). At this site, the maximum salinity was 42 g L 1, and methane hydrate was inferred to be present and stable during the last 120 ky (Stotler et al., 2009, 2010b). We caution that B, Nd, and Sr isotopic data from this site (Ne´grel et al., 2004; Casanova et al., 2005; Ne´grel and Casanova, 2005) were predominately collected in permafrost fluids, which were heavily contaminated due to mining activities (Ruskeeniemi et al., 2004; Stotler, 2008; Stotler et al., 2009). Numerous investigations have considered the origin and evolution of deep groundwaters in the Canadian and Fennoscandian Shields (Table 1); several focused directly on the potential influence of Pleistocene glaciations to deep groundwaters (Herut et al., 1990; Bein and Arad, 1992; Douglas et al., 2000; Clark et al., 2000; Starinsky and Katz, 2003). From these investigations, a hypothesis was developed suggesting that brines in the Canadian and Fennoscandian Shields formed from freezing of seawater in proglacial lagoons along the glacial margin, with subsequent recharge of these brines beneath the ice sheet (Starinsky and Katz, 2003). This hypothesis did not consider the effects of permafrost or methane hydrate formation on subsurface fluid evolution. The purpose of this study is to investigate the effects that deep permafrost formation and dissipation had on fluids in the Canadian and Fennoscandian Shields during the Pleistocene glacial/interglacial cycle. A comprehensive review of the Canadian Shield groundwater geochemical dataset and four well characterized sites from the Fennoscandian Shield will be utilized to evaluate the potential influence of freez-

47

ing-related processes to shield fluids. The geochemical signatures of permafrost and methane hydrate formation are compared with those of glacial recharge, to better understand how these processes might affect glacial recharge models in crystalline shields. Constraining the role of the glacial/interglacial cycle and how glaciation, permafrost and methane hydrate affect deep flow-system evolution, fluid movement, and the chemical evolution of groundwaters in Shield environments will allow for a direct examination and evaluation of the validity of the “cryogenic glacial lagoon shield brine origin hypothesis” presented by Starinsky and Katz (2003). 2. GEOCHEMISTRY OF FREEZING The geochemical similarities between a theoretical glacial meltwater and in situ “freeze-out” processes (the process by which solutes concentrate in residual fluids and isotopes fractionate between ice and residual fluids during freezing), such as permafrost or methane hydrate formation, are illustrated below through a review of freezing experiments designed to simulate ice formation, in situ permafrost formation, and methane hydrate formation. Data from “grey-literature” fluid freezing experiments are collated as electronic annex material accompanying this article on-line. During ice formation, distinct geochemical and isotopic patterns develop in both the residual fluid and the ice. Certain aspects of these patterns have been utilized to identify glacially recharged fluids (e.g. Siegel and Mandel, 1984; Siegel, 1991; Clark et al., 2000). The process of formation and melting of glacial ice is geochemically similar to the processes of permafrost ice and gas hydrate formation and dissipation; however, these processes were ignored when glacially recharged fluids were identified in past studies. 2.1. Stable isotopes Stable isotopes (2H/1H, 18O16O) fractionate between water, ice, and clathrates (gas hydrates) during phase changes (e.g. Craig and Hom, 1968; Hesse and Harrison, 1981; Souchez and Jouzel, 1984; Maekawa and Imai, 2000; Hesse, 2003; Maekawa, 2004). As observed in many different fluids in conservative systems (including precipitation, freshwater, seawater, and fluids of differing ionic compositions), the first ice that forms has substantially higher d2H and d18O values than the original fluid, but becomes progressively lighter as freezing progresses, until the final ice and residual fluid both have much lower d2H and d18O values than the original fluid. Regardless of initial ionic composition, d2H and d18O ratios of ice and residual fluids fall along lines with a slope between 4.32 and 7.44 (Fig. 3a, Souchez and Jouzel, 1984; Zhang and Frape, 2003), distinct from that of Craig’s (1961) global meteoric waterline (GMWL) [slope = 8.0]. Several researchers have assumed that glacial meltwater was the only fluid with a depleted stable isotope value, and utilized this property to identify glacial meltwater recharge to groundwater in various locations (Perry et al., 1982; Siegel and Mandel, 1984; Clark et al., 2000; Douglas et al.,

Site

Permafrost? depth (mbgs)

Nunavut, Canadian Shield Lupin mine Continuous, 540

Continuous

Meadowbank

Continuous

Northwest Territories, Canadian Shield Diavik Continuous Yellowknife (Con mine)

Extensive Discontinuous

Yellowknife (Giant mine)

Extensive Discontinuous

Manitoba, Canadian Shield Thompson Sporadic discontinuous

Lynn Lake (SGM Fox mine)

Sporadic discontinuous

Rock types

Structural Province/ greenstone belt

Sample depth (m)

Water types

TDS (g L 1)

Referencesc

Metasediments, banded iron formation, quartzite, phyllite, amphibolite, Au VMS, Metasediments, Cu, Zn Au

Slave

<1300

Ca–Na–Cl, Na–Ca–Cl

8–40

Ruskeeniemi et al. (2002, 2004), Frape et al. (2004), Stotler et al. (2009, 2010a, 2011) and Onstott et al. (2009)

Slave/High Lake GB

45, 430

Ca–Mg–SO4, Ca–Clb

111– 169

Na–SO4, Ca–Mg–SO4, Ca–Mg–Na–HCO3– Cl, Ca–Mg–Cl

0.15–0.79

Golder (2004)

Slave

44–404

0.13–0.56

Blowes and Logsdon (1998)

Yellowknife GB/Slave Province

<1600

Na–Mg–Ca–HCO3, Na–Mg–Ca–HCO3–Cl, Na–Mg–Cl, Na–Mg–Ca–Cl Ca–Na–HCO3, Na–Ca–Cl (d)

0.7–314

Au

Slave

30–610

Ca–SO4–Cl–HCO3, Ca–Mg–SO4–HCO3, Ca–Mg–SO4, Ca–Na–SO4, Ca–Mg–Na– SO4–HCO3, Ca–Mg–Na–SO4, Ca–Na– SO4–HCO3, Ca–Na–SO4–HCO3, Ca–Na– SO4–Cl, Na–Ca–SO4–Cl–HCO3, Ca–Na–Cl

0.7–80

Frape and Fritz (1981, 1987), Fritz and Frape (1982), Frape et al. (1984, 2004), McNutt et al. (1984), MacDonald (1986), Kaufmann et al. (1987), Fritz et al. (1987, 1994), Sherwood-Lollar et al. (1988, 1993a), Bryant (1995), Douglas (1997), Bottomley et al. (1999, 2002, 2005), Douglas et al. (2000), Clark et al. (2000), Battye (2002), Bottomley and Clark (2004), ShouakarStash et al. (2005), Greene (2005), Greene et al. (2008) and Stotler et al. (2010a) Frape and Fritz (1981, 1987), Frape et al. (1984), MacDonald (1986), SherwoodLollar et al. (1988, 1993a), Fritz et al. (1994) and Stotler et al. (2010a)

Quartzites, granite gneiss, schist, peridotite, Ni

Thompson Nickel GB

100– 1500

Na–HCO3, Na–SO4–HCO3–Cl, Na–HCO3– Cl, Mg–HCO3, Ca–SO4, Ca–Na–Cl, Ca–Cl (d)

0.5–325

61–792

Ca–HCO3, Ca–Mg–SO4, Mg–SO4, Ca–SO4, Ca–Cl

1.2–21.5

Granite, kimberlite Metabasalts in granodiorite, Au

Gartner Lee (2005, 2006a,b), Pfiffner et al. (2008) and Stotler et al. (2011)

Frape and Fritz (1981, 1987), Fritz and Frape (1982), Frape et al. (1984, 2004), McNutt et al. (1984), Kaufmann et al. (1987), Sherwood-Lollar et al. (1988, 1993a), Fritz et al. (1994), Bottomley et al. (1994), Bryant (1995), Bottomley and Clark (2004) and Stotler et al. (2010a) Frape and Fritz (1981), Fritz et al. (1994) and Stotler et al. (2010a)

R.L. Stotler et al. / Geochimica et Cosmochimica Acta 76 (2012) 45–67

High Lake

48

Table 1 Host rock types and ages, water chemistry, and groundwater geochemical data sources for studies from the Canadian Shield.

Lac du Bonnet (Whiteshell–URL)

None at present

Michigan, Canadian Shield Keweenaw (Centennial None at mine) present Ontario, Canadian Shield Red Lake (Campbell, None at Dickenson) present

Granitic gneiss

Superior

Na–Ca–Cl–SO4

0.22–33

Gascoyne et al. (1987, 1989), Purdy (1989) and Frape et al. (2004)

2100– 2800

Ca–Cl

326–406

Kelly et al. (1986)

Birch-Uchi Lake GB

640, 660,

Mg–Ca–Na–SO4–Cl,a Ca–Na–Cl

0.8,a 190

MacDonald (1986), Bottomley et al. (1994), Fritz et al. (1987) and Sherwood-Lollar et al. (1988, 1993b)

Superior

4–1050

0.03–40

Dicken et al. (1984), Gascoyne et al. (1987), Franklyn (1987), Franklyn et al. (1991) and Stotler et al. (2010a)

381– 960

Ca–Cl, Ca–SO4, Ca–Cl–SO4, Ca–HCO3, Ca–Na–HCO3, Na–HCO3, Mg–HCO3, Ca– Mg–HCO3, Ca–Na–Mg–HCO3, Ca–Na–Cl, Ca–Na–SO4–HCO3–Cl, Ca–Na–HCO3–Cl Na–Ca–Cl, Ca–Cl, Ca–Na–SO4,a Ca–Na– Cl

2.1–50

MacDonald (1986), Fritz et al. (1987), Sherwood-Lollar et al. (1988, 1993b), Bottomley et al. (1994), Stotler (2008) and Stotler et al. (2010a)

0–836

Ca–HCO3 (s) mixing to Na–Cl (d)

0.08–4.3

Kamineni (1987), McNutt et al. (1987), Gascoyne et al. (1987), Bottomley et al. (1990, 1994) and Frape et al. (2004) Frape and Fritz (1982), Fritz and Frape (1982), Frape et al. (1984, 2004), McNutt et al. (1984); Frape and Fritz (1987), Kaufmann et al. (1987), Sherwood-Lollar et al. (1988, 1993b), Bottomley et al. (1994), Fritz et al. (1994), Bryant (1995), Bottomley and Clark (2004), Shouakar-Stash et al. (2005) and Stotler et al. (2010a) Frape and Fritz (1982, 1987), Frape et al. (1984), McNutt et al. (1984), SherwoodLollar et al. (1988, 1993b, 2006), Bottomley et al. (1994), Doig (1994), Fritz et al. (1994), Montgomery (1994), Bryant (1995) and Stotler et al. (2010a) Frape and Fritz (1981), Fritz and Frape (1982), MacDonald (1986) and Fritz et al. (1994) MacDonald (1986) and Bottomley et al. (1994)

Basalts, tholeitic, conglomerates Mafic volcanics, ultramafics, metavolcanics, metasediments, granitic plutons GranodioriteGranitic pluton

Atikokan (Eye-Dashwa)

None at present

Elliot Lake (Panel, Stanleigh, Dennison)

None at present

East Bull Lake (EBL)

None at present

Huronian metasediments (conglomerate, quartzite, argillite), U Gabbro and anorthosite

Sudbury – North (Creighton, Fraser, Garson, Lockerby, North, Onaping, Strathcona, Victor, Dowling Wells)

None at present

Mafric intrusive (Norite)

Superior

152– 1600

Ca–Na–SO4–HCO3, Ca–Na–SO4–Cl– HCO3, Na–Cl–HCO3, Mg–Ca–SO4, Na–Cl, Ca–Na–Cl, Ca–Cl

0.11–254

Sudbury – South (Copper Cliff South, Copper Cliff North)

None at present

Mafric intrusive (Norite)

Superior

121– 1219

Ca–Cl, Ca–Mg–SO4, Ca–Na–Cl–SO4

0.26–257

South Bay (Selco)

None at present

Felsic volcanics

Superior

15–595

0.05–7.9

Kirkland Lake (Macassa, Lakeshore)

None at present

Conglomerate, greywacke, tuff, syenite

Superior/ Abitibi GB

61– 1966

Ca–Na–HCO3, –Cl, Ca–HCO3, Ca–HCO3– SO4, Ca–SO4 (0–320 m) Ca–Na–SO4, Ca– Na–Cl (595 m) Ca–Mg–Na–SO4, Ca–Mg–HCO3, Ca–Mg– HCO3–SO4, Ca–Mg–Na–HCO3, Ca–Mg– Na–HCO3–SO4 nitrate

0.5–14.8

Freezing processes and evolution of Shield groundwater

2.1– 1098

(continued on next page)

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50

Tabel 1 (continued) Permafrost? depth (mbgs)

Rock types

Structural Province/ greenstone belt

Sample depth (m)

Water types

TDS (g L 1)

Referencesc

Timmins (Kidd Creek)

None at present

Mafic and felsic metavolcanics, diorite

Superior/ Abitibi GB

610,a 1402– 1616

Na–Ca–Cl,a Ca–Cl

1.4,a 88–129

Chalk River

None at present

Metasedimentary gneiss, qtz. Monzonite, gneiss

Grenville

341– 576

Na–Cl, Na–HCO3

0.2–0.7

Bottomley et al. (1994), Doig (1994), Montgomery (1994), Bryant (1995), Westgate (1998) and Sherwood-Lollar et al. (2002) Gascoyne et al. (1987)

Quebec, Canadian Shield Chibougamau (Copper Rand mine)

Isolated

Ca–Na–Cl, Ca–Cl

326–406

Guha and Kanwar (1987)

Isolated

Superior/ ChibougamauAbitibi GB Superior/ Abitibi GB

640, 869

Matagami/Norita

Metaanorthosite, Cu– Au Mafic volcanic and ultramafics

<670

Shallow: Ca–Na–Mg–HCO3, Ca–Mg–SO4, Ca–SO4 Deep: Ca–Cl, Ca–Na–Cl

0.5–240

Selbaie

Isolated

Superior/ Abitibi GB

60–240

Ca–Mg–HCO3, Mg–Ca–HCO3

0.4

Val d’Or (Sigma, Lamaque)

None at present

Felsicintermediate volcanics, metasediments overly, Zn–Cu– Ag Mafic – felsic volcanics

Frape and Fritz (1987), Jones (1987), Sherwood-Lollar et al. (1988, 1993a,b), Fritz et al. (1994), Frape et al. (2004) and Stotler et al. (2010a) MacDonald (1986) and Fritz et al. (1994)

Superior/ Abitibi GB

73, 116, 1737

Ca–HCO3–SO4 (s), Ca–Cl (d)

0.4–66

Sweden, Fennoscandian Shield ¨ spo¨ A None at present

Granite, diorite, greenstone

Finland, Fennoscandian Shield Olkiluoto None at present

Palmottu Ha¨stholmen

a b c

None at present None at present

Migmatitic felsic gneisses, pegmatitic granites Mica gneiss, pegmatite Rapakivi granite

MacDonald (1986), Fritz et al. (1987), Sherwood-Lollar et al. (1988), Fritz et al. (1994) and Stotler et al. (2010a) Laaksoharju and Rehn (1999), Laaksoharju et al. (1999, 2008) and Ne´grel et al. (2005)

Southern Finland Sedimentary– volcanic Complex Uusimaa Belt Wiborg Rapakivi Batholite

10–930

Shallow: Ca–Na–Mg–HCO3, Na–Ca–Cl– SO4 (<100 m), Na–Cl–SO4, Na–Cl (100– 300 m), Na–Ca–Cl, Ca–Na–Cl

0.3–83

Pitka¨nen et al. (1996, 1999, 2004), Ne´grel et al. (2005) and Pitka¨nen and Partamies (2007)

<407

Shallow: Ca–HCO3, NaHCO3, NaSO4 Deep: Na–Cl Shallow: Ca–Mg–Na–HCO3, Na–Ca–Mg– Cl–HCO3–SO4 (<50 m), Na–Ca–Mg–Cl– SO4 (50–600 m), Na–Ca–Cl

0.1–1.6

Blomqvist et al. (2000) and Ne´grel et al. (2003, 2005) Pitka¨nen et al. (2001) and Ne´grel et al. (2005)

5–900

Bottomley et al. (1994) significantly lower TDS than previous studies. Samples contaminated by drilling salt. d37Cl values reported in Kaufmann et al. (1987) have been shown to be measured incorrectly by Bryant (1995).

0.2–32

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Site

Freezing processes and evolution of Shield groundwater

51

Fig. 3. Relationship between d2H and d18O (a) during freezing of shield groundwaters; (b) Canadian Shield sites located in continuous permafrost; (c) sites located in discontinuous, sporadic, and isolated permafrost; (d) Canadian Shield sites not currently located in permafrost; (e) Fennoscandian Shield sites; (f) brines only. Note during freezing, the first ice is isotopically heavier than the initial fluids, with residual fluids significantly lighter. The slope of d2H and d18O during fluid freezing evolution slightly less than the GMWL. The effect of mining is conceptualized in panels (g) original system and (h) disturbed conditions. Data are from various sources cited in Table 1. Due to the complexity of the diagram, sites are displayed individually in accompanying Supplementary online material.

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2000; McIntosh et al., 2002; McIntosh and Walter, 2005, 2006; Person et al., 2007). However, similar depleted stable isotope values may result from in situ ice formation during permafrost growth (e.g. Mackay and Lavkulich, 1974), and isotopic fractionation factors (for both stable hydrogen and oxygen isotopes) are the same between the ice–water pair and the clathrate–water pair: 1.0178–1.0206 (a2Hice–water) and 1.014–1.022 (a2Hclathrate–water); 1.0027–1.0032 (a18Oice– 18 water) and 1.00268–1.0037 (a Oclathrate–water) (Craig and Hom, 1968; O’Neil, 1968; Suzuoki and Kimura, 1973; Mackay and Lavkulich, 1974; Trofimuk et al., 1974; Hesse and Harrison, 1981; Davidson et al., 1983; Handa, 1984; Hesse et al., 1985; Ussler and Paull, 1995, 2001; Matsumoto, 1998; Matsumoto and Borowski, 2000; Maekawa and Imai, 2000; Hesse, 2003; Zhang and Frape, 2003; Maekawa, 2004). Thus, in areas where two (or more) of these freezing-related processes may have affected fluid evolution, the stable isotopic composition represents a non-unique solution for distinguishing among the three freezing processes. Despite this, stable isotopic compositions of fluids are clearly an important tool for determining whether or not fluid has been affected by one of the three freezing processes.

2.2. Solute exclusion During freezing, solutes are excluded from ice and hydrate phases, concentrating into the residual fluids (e.g. Mackay and Lavkulich, 1974; Konrad and McCammon, 1990; Ussler and Paull, 1995). As a result, glacial meltwater has extremely low dissolved solids, and most solutes in recharged glacial meltwater should have been leached from the bedrock (Douglas et al., 2000; Clark et al., 2000). Thus, in an area such as the Canadian Shield, a three end-member mixing diagram between d18O and Cl may be constructed (Fig. 4a), with (1) a modern, fresh meteoric recharge component (displayed in the diagram as mean modern Yellowknife, NWT, Canada, precipitation), (2) an initially low-salinity glacial recharge component, with the saline component leached from bedrock after glacial fluid infiltration, and (3) geologically old resident Shield brine (Douglas et al., 2000; Clark et al., 2000). It should be noted that near coastal sites in the Fennoscandian Shield, a fourth endmember derived from the palaeo-Baltic Sea, has also been identified (Pitka¨nen et al., 2001, 2004). The evolution of this end-member will be discussed later in the paper. Residual fluid solute concentrations resulting from the in situ freezing processes are direct evolutionary products of the original fluid. Thus, during in situ freezing, d18O and Cl co-vary and result in an end-member essentially identical to a hypothetical intruded glacial-meltwater recharge component (with salinity leached from bedrock) (Fig. 4a). Despite the similar result, the evolutionary pathway is important; for glacial-meltwater recharge, the initial fresh fluid already has a low d18O value, but for in situ freezing, the initial fresh fluid has a much higher initial d18O value. Such a distinction is evident in the results of the “freeze-out” experiments conducted by Zhang and Frape (2003); during freezing a nearly linear relationship develops between d18O and the log of chloride concentration for each

of the fluids with differing solute compositions. As d18O ratios decrease, chloride concentrates in the residual fluid, regardless of the initial fluid solute composition (Fig. 4a). Thus simply adding chloride concentration data to stable isotope values allows fluids that have been affected by one of the three freezing processes (glacial recharge with saline component leached from bedrock, in situ permafrost formation, in situ methane hydrate formation) to be identified. In scenarios with a semi-log-linear d18O–Cl relationship similar to those observed in the experiments by Zhang and Frape (2003), in situ processes can be identified. Fluids with an oxygen isotopic value less than modern meteoric precipitation, and increasing salinity with no change in oxygen isotopic value, as described by Douglas et al. (2000) and Clark et al. (2000), indicate glacial recharge. However in scenarios with significant mixing, the three freezing processes remain indistinguishable. As solutes concentrate and temperatures decrease during fluid freezing, minerals may precipitate, further altering the fluid composition (e.g. Marion and Farren, 1999; Marion et al., 1999). Because of this, it has been proposed that a comparison of Na/Cl and Br/Cl ratios should differentiate between water affected by freezing and evaporative processes (Herut et al., 1990). As freezing progresses, ionic Na/Cl and Br/Cl ratios respond differently depending on the precipitating mineral phases, which are dependent on the fluid’s initial solute composition (Fig. 5a). As a result, Na/Cl ratios vary differently for Na- and Ca-dominated fluids with increasing residual fluid concentration (Fig. 5a). Very little change in the Na/Cl or Br/Cl ratios occurs during freezing for Na-dominated fluids of different chemical composition, including seawater (Fig. 5a). However, Na/ Cl and Br/Cl ratios generally increase with increasing residual fluid concentration in Ca-dominated fluids (Fig. 5a). Neither Na/Cl nor Br/Cl ratios were significantly affected by differing anion dominance (Cl vs. SO4) (Fig. 5a). Thus, when the resulting trends are plotted on a Na/Cl vs. Br/Cl graph (Fig. 5a), Na-dominated fluids have a nearvertical slope, Ca-dominated fluids have a slightly positive slope, and fluids with neither Na- nor Ca-dominance could be expected to plot along a slope between pure Na- and Ca-dominant fluids (regardless of the initial anion composition). Several important observations should be noted. As discussed previously by Herut et al. (1990), fluid chemical evolution due to freezing is distinct from seawater evaporation (Fig. 5a). However, data are not available to evaluate Br/Cl and Na/Cl ratio evolution during other water–rock interaction processes, such as silicate hydration, that commonly affect fluids in crystalline rock. It has further been suggested that any correlation observed comparing these ratios could be the result of an induced correlation (Lenahan et al., 2011). With this in mind, at this time, investigation of geochemical ratios in crystalline fluids may point toward freezing or evaporation as a potentially significant geochemical process, but cannot be considered conclusive without supporting isotopic data. Importantly, it is noted that no experimental data support the hypothesized (and extrapolated) seawater freezing trend described by Herut et al. (1990) and Starinsky and Katz (2003) (Fig. 5a).

Freezing processes and evolution of Shield groundwater

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Fig. 4. Relationship between d18O and Cl (a) during freezing; groundwaters sampled from Canadian Shield sites located in (b) continuous permafrost regions; (c) in discontinuous, sporadic, and isolated permafrost regions; (d) sites not located in or near permafrost; (e) Fennoscandian Shield sites. Arrows indicate progression during freezing. Shield groundwater freezing data from Zhang and Frape (2003). The glacial three end-member mixing diagram modified from Douglas et al. (2000) and Clark et al. (2000), compared with an in situ end-member three end-member mixing scenario shows permafrost and intruded glacial meltwater end-members are similar. Data are from various sources cited in Table 1. Due to the complexity of the diagram, sites are displayed individually in accompanying Supplementary online material.

3. CRYSTALLINE SHIELD FLUIDS With the experimental background in mind, field data from the Canadian and Fennoscandian Shields are reviewed to determine how the freezing processes influenced geochemical evolution. Samples collected over the last 30 years from 39 sites at 24 locations across the Canadian

Shield, and four sites from the Fennoscandian Shield are compiled for this analysis (Table 1 and Fig. 1). Eleven Canadian sites are located within various permafrost zones; the 15 remaining sites are not currently located within or near permafrost areas. Of the 12 permafrost sites (Fig. 1), High Lake, Lupin, Lac de Gras (Diavik), and Meadowbank Gold are the only sites located within continuous

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Fig. 5. Na/Cl vs. Br/Cl ratios from (a) freezing experiments, (b) Canadian Shield sites located in continuous permafrost, (c) sites located in discontinuous, sporadic, and isolated permafrost; (d) Canadian Shield sites not currently located in permafrost; (e) Fennoscandian Shield sites; (f) brines only. The seawater evaporation trend (McCaffrey et al., 1987), frozen seawater and predicted seawater freezing curve (Herut et al., 1990) are included for reference. Data are from various sources cited in Table 1. Due to the complexity of the diagram, sites are displayed individually in accompanying Supplementary online material.

permafrost; Con and Giant are within the zone of extensive discontinuous permafrost; Thompson and Lynn Lake are located in areas of sporadic discontinuous permafrost; and Copper Rand (Chibougamau), Matagami-Norita, and Selbaie are in or very near the zone of isolated permafrost. Although only northern portions of the Canadian Shield are currently covered by permafrost, climatic and surface conditions varied dramatically throughout the Pliocene

and Pleistocene. In the southern Canadian Shield, a shallow permafrost regime likely formed (Fig. 2) in advance of, and following, the periodic continental glaciations that covered the entire Canadian Shield (Fig. 1). Therefore, data from the remaining 13 Canadian Shield sites are considered for this analysis as well. Finally, four well-studied Fennoscandian Shield sites are also examined; three coastal (Olkiluoto, Ha¨stholmen, ¨ spo¨) and one located inland (Palmottu). Glaciers and A

Freezing processes and evolution of Shield groundwater

permafrost are currently absent from both Finland and Sweden; however, the entire Fennoscandian Shield was glaciated at the last glacial maximum (e.g. Peltier, 1994; Kleman and Ha¨ttestrand, 1999). Although permafrost features occur only sporadically, permafrost can be considered similarly forming in front of advancing and retreating glaciers (Saarnisto, 1977; Kukkonen and Sˇafanda, 2001; Kukkonen and Jo˜eleht, 2003; Pitka¨ranta, 2009). The Diavik data (reported by Blowes and Logsdon, 1998, available with the Supplementary online information accompanying this article) shows that geochemical trends related to an in situ freezing process are found in crystalline shields outside of controlled laboratory conditions. The samples were collected in permafrost during tunnel excavation prior to mining operations, minimizing the potential for mine dewatering to affect natural flow conditions. Thus, these samples provide an excellent view of geochemical conditions within crystalline rock permafrost settings. Considerable scatter occurred in ionic data, with no dominant cation, yet general trends were still observed when Na/Cl and Br/Cl ratios were compared (Fig. 5b). The most dilute fluids followed the Na-dominated fluid freezing trend, but as dissolved solids increased (0.30–0.56 g L 1), the Cadominated freezing trend was observed, although the most concentrated fluids had lower Na/Cl and Br/Cl ratios, opposite from freezing experiment results (Fig. 5b). The relationships between the stable isotope (d2H, d18O) compositions (Fig. 3b) and Cl and d18O (Fig. 4b) values of Diavik water samples were also consistent with trends observed in freezing experiments. The semi-log linear relationship between d18O and Cl (Fig. 4b) indicates these fluids were derived from a single in situ freezing process, rather than mixing between a fresh meteoric end-member and a glacially derived fluid (with salinity leached from the rock). Geochemical samples from most other Canadian Shield sites were collected up to 60 years after excavation commenced, with significant disruption to natural hydraulic flow. At these sites, fluid mixing (and sometimes contamination) is readily apparent; thus distinguishing between fluid freezing processes is more complicated. Based on the relationship of Cl and d18O, geochemical evidence of in situ freezing was found at six of the other 10 permafrost sites (no trend observed at Matagami/Norita; not enough data from Selbaie, Meadowbank Gold and Copper Rand) and at five of the 12 non-permafrost sites, including the Lac du Bonnet-URL site, South Bay Mine, East Bull Lake, Elliot Lake, and the Eye-Dashwa Pluton (Fig. 4b–d). Data from many of the Canadian Shield sites currently located in permafrost indicate mixing and contamination affected many of the sampled fluids with salinities less than 50 g L 1. Of the non-permafrost Canadian Shield sites, only groundwaters sampled from the Lac du Bonnet URL site follow all three of the geochemical trends observed in the freeze-out experiments. Elliot Lake, East Bull Lake, and South Bay groundwaters have d18O signatures similar to those that would be expected for waters influenced by a freezing process, but Na/Cl vs. Br/Cl ratio patterns were scattered and did not follow general freezing experiment trends (Fig. 5d). Fluids at other sites, such as Val d’Or, had Na/Cl vs. Br/Cl ratio patterns similar to

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freeze-out experiments (Fig. 5d), but d18O values were not similar (Fig. 4d). Some characteristics of freezing were also observed in Eye-Dashwa fluids, but the data were highly scattered. Thus, some fresh, brackish and saline waters at sites not currently located in or near permafrost did have geochemical characteristics indicative of glacial/interglacial influence to groundwater geochemistry, but the evidence was equivocal. At some sites (e.g. Selbaie, Copper Rand, Centennial Mine, Chalk River, High Lake, Meadowbank Gold), not enough samples were collected, and determining whether fluid freezing affected the chemical evolution was not possible. Including sites without enough data, geochemical evidence of freezing was observed at 59% of permafrost sites and 29% of non-permafrost sites; however when the population included only sites where clear determinations were possible, these values changed to 87% and 84%, respectively. Finally, the most concentrated shield fluids form a geochemical end-member distinct from freezing affected fluids (Figs. 3f, 4b–e, 5f). These concentrated fluids will be discussed further in the next section. Fennoscandian Shield fluids have also been affected by ¨ spo¨, Ha¨sfreezing processes. At the three coastal sites (A tholmen and Olkiluoto), the Baltic Sea/Littorina Sea (Baltic sea stage 6000–8000 ybp), formed a fourth end-member for mixing scenarios (referred to earlier in the paper), resulting from the submergence of these coastal sites below sea-level between 8000 and 3000 ybp (Fig. 4e). Fluids from these three coastal sites reflect mixing between all four end-members (Littorina, modern precipitation, relic shield brine and a paleo-fluid with depleted d2H–d18O). At all Fennoscandian sites, fluids with the most depleted signatures are found at depths of 200–400 m (this was also the maximum sampling depth at Palmottu). The mixing of Littorina Sea derived water partly masks the cold climate signature at the coastal sites due to its heavy isotope composition. The cold climate signature is most prominent at Palmottu, which is situated at a notable Salpausselka¨ III ice marginal formation (end moraine?). Due to slightly cooling climate, the retreat of the continental ice sheet stopped temporarily at the Palmottu area (for hundreds of years) producing significant amounts of meltwater and building extensive glaciofluvial deposits (Ma¨kinen and Palmu, 2008). 3.1. Shield brines – surficial freezing derived? The brine component (TDS >100 g L 1) of Canadian and Fennoscandian Shield fluids have been suggested to have formed during Plio-Pliestocene glaciations, resulting from surficial freezing of seawater along ice-sheet margins with post-genetic density induced emplacement (Fig. 6A and B, Starinsky and Katz, 2003). However, evidence suggests at least some Canadian Shield brines are too old to support such a hypothesis (Bottomley et al., 2002, 2003). Additional physical and geochemical evidence, described below, provide further arguments against a surficial seawater freezing origin for shield brines. Evidence and processes, from subglacial flow, permeability, distribution of saline glacial margin lakes, Na/Cl and Br/Cl ratios, stable isotope ratios and thermodynamic equilibrium models, all suggest cryogenic freezing of seawater along glacial and continental

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Fig. 6. Comparison of flow under glacial ice between (A, B) Starinsky and Katz (2003) hypothesis, and (C, D) a conceptual model that incorporates subglacial pressures and flow.

margins was not the source of Canadian and Fennoscandian Shield brines. Starinsky and Katz (2003) postulation ignores important physical processes, including hydraulic gradient and permeability. First, the hypothesis suggested salinity formed on continental margins, indicating a continental scale reversal of groundwater flow. Flow reversal during glacial periods has been identified; however, this resulted from increased pressures beneath the glacial ice mass changing subsurface pressure fields, resulting in groundwater flow outward from beneath the ice sheets (Boulton et al., 1995; Grasby et al., 2000; Person et al., 2007; Lemieux et al., 2008a). Recent simulations of continental scale flow and recharge, incorporating density-dependent flow, hydromechanical loading, subglacial infiltration, isostasy, and permafrost development suggested groundwater recharge typically occurs during glacial advance, with discharge occurring during glacial retreat, a result of reduced pressures when the weight of the ice mass was removed (Fig. 6C and D). Evidence of hydraulic gradients required for oceanic fluid infiltration through the forebulge has not been identified. Thus, flow towards the center of the continent as required by the Starinsky and Katz (2003) hypothesis is not physically supported. The hypothesis also ignores the low permeability of shield rocks. For instance, brines at the Copper Rand site were the most concentrated fluids identified, with isotopic and geochemical characteristics similar to other shield brines. The brines were located in vugs and were not part of an active flow system (Guha and Kanwar, 1987). Using a simple, “back of the envelope” groundwater flow calculation, given an average hydraulic conductivity for intact rock of 10 12 m s 1 (Freeze and Cherry, 1979) and a generous hydraulic gradient of 1.0 in the area surrounding the vug, fluid could be expected to travel only 31.5 m in 100,000 years. This calculation indicates barely enough time elapsed during a glacial cycle for fluid to travel to

the vug from a nearby fracture, let alone from the edge of the continent, 1100 km away (Fig. 1). Thus, permeability within crystalline shields is too low for brine transport from continental margins over the time scale proposed by Starinsky and Katz (2003). Characteristics of proglacial environments are also dissimilar from those proposed by Starinsky and Katz (2003). Today, saline lakes along the glacial margin are the exception. A significant majority of lakes in Greenland, the Arctic and Antarctic, including along glacial fringes, are composed of freshwater (Douglas and Smol, 1994; Ru¨hland and Smol, 1998; Anderson et al., 2001; Hamilton et al., 2001; Lim et al., 2001, 2005; Michelutti et al., 2002; Lim and Douglas, 2003; Keatley et al., 2007; Kumke et al., 2007). The few instances of saline lakes found near glacial margins (in the McMurdo dry valleys of Antarctica and in Greenland) have salinity derived from multiple processes including marine aerosols, glacial and chemical weathering, ancient marine waters and evaporation, and are not limited to cryogenic concentration (Anderson et al., 2001; Willemse et al., 2004; Lyons et al., 2005). Finally, in continental shelf regions where proglacial lakes formed during the last glacial advance, freshwater aquifers are found in disequilibrium with overlying ocean water (Vaikma¨e et al., 2001; Person et al., 2006). Following the Starinsky and Katz (2003) hypothesis, either concentrated brines or ocean water would occur in these regions, rather than freshwater. Stable isotope data (d2H, d18O) indicate shield brines did not evolve through a freezing process. Unlike most other groundwaters, shield brines plot above the GMWL for d2H and d18O (Fig. 3f), a result of water–rock interaction and silicate hydration (Fritz and Frape, 1982; Frape et al., 1984; Pearson, 1987; Longstaffe, 2000; Ziegler and Longstaffe, 2000). Fennoscandian fluids in d18O–Cl plots clearly show a Baltic or paleo-Baltic seawater end-member as a unique end-member, and not the origin of brines (Fig. 4e). The lack of trends toward such an end-member

Freezing processes and evolution of Shield groundwater

from other shield fluids suggests that seawater was not a significant factor in the evolution of shield brines, i.e., other sources or processes had a stronger influence. In fact, the Littorina end-member observed at the three coastal Fennoscandian sites actually precludes a cryogenic sea-water origin for the saline fluids. If seawater were to cryogenically concentrate, a semi-log straight line relationship would form from the Littorina Sea end-member towards increasing salinity and decreasing d18O values, similar to the freezing experiments (Fig. 4a). This was not observed – rather, as fluids at these sites “evolved” from the Littorina endmember towards decreasing d18O values, salinity actually decreased – indicating the Littorina end-member mixed with fluids at the sites but did not evolve along a cryogenic pathway. As such a process was not observed at sites located next to a large saline sea in the Fennoscandian Shield, such a process is also not expected to affect fluid evolution in the center of the Canadian Shield, at significant distance from the nearest sea or ocean. Shield brines do not have a stable isotopic signature suggestive of seawater, evaporation, freezing, or dissolution of evaporite minerals, indicating the majority of brines and highly saline shield fluids were not formed or significantly affected by either freezing of, or evaporation of, seawater. Further, the most concentrated shield fluids are a distinct d18O–Cl end-member that could not have formed during in situ freezing of fluids (e.g. permafrost or methane-hydrate formation), surficial freezing of seawater, evaporation of seawater or dissolution of evaporites. The relationships between Na/Cl and Br/Cl ratios also do not support a freezing evolutionary pathway for the brines. Shield brines have unique Na/Cl and Br/Cl ratios (Fig. 5f) compared with both experimental evaporation of seawater and freezing of modern-day shield fluids or seawater (McCaffrey et al., 1987; Herut et al., 1990; Zhang and Frape, 2003), indicating the brines were not derived from a freezing process. While some shield brines do fall near the hypothesized seawater freezing trend presented by Herut et al. (1990) (Fig. 5a), Herut et al. extrapolated the trend significantly beyond the final experimental data, and along a completely different slope than indicated by the experimental data; further this trend was not reproduced during other freezing experiments that considered Na–Cl type fluids (Zhang and Frape, 2003). Finally, the apparent basis for the Herut et al. (1990) extrapolation were Br and Cl data from Nelson and Thompson (1954); however Br data are not presented by Nelson and Thompson (1954), as cited by Herut et al. (1990). Thus, the hypothesized seawater freezing trend presented by Herut et al. (1990) is not currently supported, and therefore the fact that some, but not most, shield brines fall along such a “trend” does not support the hypothesis. In the surficial cryogenic concentration of seawater hypothesis, Starinsky and Katz (2003) suggest that shield brine Sr2+ has a different origin than Ca2+. In this postulation, Ca2+ is derived from concentration of seawater through freezing and Sr2+ is derived from water–rock interaction (Starinsky and Katz, 2003). However, co-evolution of Sr2+ and Ca2+ in Canadian Shield groundwaters is well-established (McNutt et al., 1984, 1987, 1990; McNutt,

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1987). If Pleistocene seawater cryogenically concentrated to form Ca-rich brines as hypothesized (Starinsky and Katz, 2003), seawater Sr2+ would also concentrate in the residual fluid, and the 87Sr/86Sr ratios (0.707-–0.755) of Canadian Shield brines should be the same as Pleistocene seawater (0.709, Capo and DePaolo, 1990). While brines from only one site (Matagami/Norita) had similar 87Sr/86Sr composition to Pleistocene seawater; d18O signatures did not indicate freezing (Fig. 7a). Further, these samples are expected to have such a 87Sr/86Sr composition as they are from Rb-poor, ultramafic rocks (McNutt et al., 1990). It has been previously noted that the large variation in 87 Sr/86Sr ratios observed across the Shields and within specific sites precludes brine migration over hundreds of kilometers; but rather supports water–rock interactions (Ne´grel and Casanova, 2005). Specific 87Sr/86Sr ratios also indicate that even if incursion of paleo-seawater (Devonian) or sedimentary brines occurred, water–rock interactions completely masked the seawater/sedimentary Sr isotopic ratio (Ne´grel and Casanova, 2005; Ne´grel et al., 2005). Nd isotopic data from at least one site, Palmottu, supports these conclusions (Ne´grel et al., 2001). We add that Shield brines do not have 87Sr/86Sr ratios similar to Pleistocene seawater, and thus 87Sr/86Sr data are not supportive of a surficial cryogenic concentration of seawater origin for shield brines during Pleistocene glaciation. Similarly, the Starinsky and Katz (2003) hypothesis indicates chlorine and bromine ions in the brine were derived from Pleistocene seawater; however this also is not supported by d37Cl and d81Br data. During freezing, 37Cl enriches in ice, but because the amount of chloride incorporated into ice is so small, the effect on the fluid reservoir is negligible when compared to brine concentrations (15–30 mg L 1 vs. 100’s g L 1, Zhang and Frape, 2003). It is currently accepted that oceanic d37Cl ratios do not vary from 0& standard mean ocean chloride (SMOC), except for small, semi-isolated basins that have terrestrial Cl inputs, such as the Black Sea, with minimal d37Cl variation in seawater over the Phanerozoic (Kaufmann, 1984; Kaufmann et al., 1984; Eggenkamp et al., 1995; Eastoe et al., 2001, 2007; Godon et al., 2004; Stewart and Spivack, 2004; Sharp et al., 2007). The bromine isotope system is less well studied, but the value of modern ocean bromide is 0& (Standard Mean Ocean Bromide – SMOB), and fractionation processes are expected to behave in a manner similar to stable chlorine isotopes (Stotler et al., 2010a). Thus if the salt source for brines in the Canadian Shield were derived from surficial freezing of seawater, both the d37Cl and d81Br values would be expected to equal the seawater value (0& SMOC, SMOB). However, even in the brine samples with the most depleted d18O values, the d37Cl and d81Br values are not consistent with a seawater origin (Fig. 7b and c). Finally, an examination of fluid freezing using thermodynamic equilibrium models suggests that the brines have not been affected by in situ freezing. Two different seawater freezing models (Fig. 8, Marion et al., 1999) suggest mirabilite (Na2SO4
10H2O) precipitation occurs between 6.3 and 8.2 °C, with hydrohalite (NaCl
2H2O) the first chloride mineral to precipitate during freezing ( 22.9 °C). Antarcticite (CaCl2
6H2O) precipitates at the eutectic of one

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Fig. 7. Relationships between d18O and (A) 87Sr/86Sr, (B) d81Br, and (C) d37Cl. The grey shaded areas indicate Pleistocene–Holocene seawater isotopic values for 87Sr/86Sr, d81Br, and d37Cl (within method detection limits). By definition, modern seawater d18O is 0&. Data are from various sources cited in Table 1. Due to the complexity of the diagram, sites are displayed individually in accompanying Supplementary online material.

Fig. 8. Comparison of seawater freezing pathways from Marion and Farren (1999). (a) Gitterman pathway (Gitterman, 1937); (b) Ringer– Nelson–Thompson pathway (Ringer, 1906a,b; Nelson and Thompson, 1954).

freezing pathway ( 53.8 °C), but MgCl2
12H2O forms at the eutectic point of another pathway ( 36.2 °C). Significantly, only mirabilite is stable within or near the lowest temperatures ( 7 to 8 °C, Natural Resources Canada, 1995) currently found beneath the seasonally affected layer of permafrost on the Canadian Shield. Even as far north as

76°N, in the Arctic Archipelago, the subsurface thermal minima ( 15 °C) are found within the upper 100 m below ground surface (mbgs) (Natural Resources Canada, 1995). Thus, temperatures have not become cold enough and have not persisted to depths great enough for an in situ freezing process to form brines.

Freezing processes and evolution of Shield groundwater

3.2. Freezing-evolved fluids – glacial? It is easy to suggest the depleted d18O values in the fresh to saline shield fluids are simply a relic of glacially induced recharge, particularly, glacial meltwater recharge. The depths at which the deepest d18O minima occur are deeper than the thickest recorded permafrost extent in the Canadian Arctic (700 m in the Arctic Islands, Natural Resources Canada, 1995). Even near the URL in the currently permafrost-free southern Canadian Shield, paleo-permafrost depths were likely only 80 m (Tarasov and Peltier (2007), see Fig. 2). At sites where fluids are suspected to have been affected by freezing, d18O becomes depleted from surface water/modern precipitation values to a minimum value (less than modern precipitation), then begins to become more enriched again deeper in the rock column. At sites currently affected by permafrost, the d18O minimum occurs between 700 and 1100 mbgs. This is most pronounced at the Con, Giant, and Thompson mines; but data from most sites are too sparse to observe such a relationship (Fig. 9a). At non-permafrost sites, the d18O minima are typically found at depths between 200 and 600 m (Fig. 9b). However, any analysis indicating fluid evolution due to a freezing process requires a clear understanding of subsurface temperature and pressure evolution during and after glaciation, fluid density effects, and flow and transport evolution at the site prior to sample collection. In the case of the Canadian Shield dataset, flow and transport evolution includes attention to the influence of excavation of mine caverns. As the subsurface freezing-evolved fluid component is most often attributed to glacial recharge, the following discussion focuses on glacial recharge and explores when glacial recharge is likely and how other processes might interfere with an assumption that all

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subsurface d18O minima in the Canadian and Fennoscandian Shields are a result of glacial recharge. Although permafrost ice has extremely low permeability (e.g. Burt and Williams, 1976), glacial ice acts as an insulating barrier, warming parts of the subsurface and melting permafrost (Hughes, 1995; Kleman and Glasser, 2007). Thus when permafrost is present, even during glacial formation, recharge is limited to pre-existing taliks (unfrozen hydraulic conduits through permafrost). Currently, no hydraulic models consider how taliks affect subsurface flow during glacial onset, although some preliminary hydraulic flow modeling indicated taliks can act as both groundwater recharge and discharge zones (Vidstrand et al., 2008). As permafrost melts, subsurface permeability increases, leading several to conclude that recharge associated with glaciers would come from glacial meltwater during glacial retreat (Clark et al., 2000; Douglas et al., 2000). However recent hydrogeologic-glacial modeling suggested that glacial-induced groundwater recharge would actually be favored earlier in the lifespan of a glacier, induced by downward hydraulic gradients created from the increased pressure provided by the ice mass (Lemieux, 2006; Lemieux et al., 2008a,b,c). Such recharge scenarios still rely on meltwater at the base of the glacier and melting of any pre-existing permafrost, regardless of hydraulic gradient. While it has been recognized that surface loading and unloading of glacial ice dramatically changes subsurface pressure and flow regimes (Grasby et al., 2000; Lemieux, 2006; Person et al., 2007; Lemieux et al., 2008a), the effects of these pressure changes on subsurface geochemistry are generally ignored. An important and common component of subsurface geochemical system at many of the Canadian and Fennoscandian sites studied is methane (SherwoodLollar et al., 1988, 1993a,b, 2002, 2006; Sherwood-Lollar,

Fig. 9. Relationship of d18O with depth in groundwaters for Canadian Shield sites located in or near permafrost areas (a) and sites not currently located within permafrost areas (b). Symbols are scaled relative to TDS, with brines the maximum size, brackish and freshwaters the smallest size, and saline waters in-between. Data are from various sources cited in Table 1. Due to the complexity of the diagram, sites are displayed individually in accompanying Supplementary online material.

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Fig. 10. Methane-hydrate and ice phase equilibria in freshwater calculated from empirical measurements discussed by Sloan (1998). The Lupin, Canada, geothermal gradient and modern hydrostatic pressures are provided as an example. (a) hydrate stability for present-day conditions (b) Hydrate stability field with a kilometerthick glacier, exerting 10,000 kPa of pressure (1 km of glacial ice), with a “warm” glacial base and a conservative geothermal gradient.

1990; Doig, 1994; Montgomery, 1994; Westgate, 1998; Pitka¨nen and Partamies, 2007; Stotler et al., 2010b). If enough methane is present in the subsurface, shifting subsurface pressure and temperature fields in crystalline shields can result in formation and dissipation of methane hydrate (Stotler et al., 2010b). Thus while permafrost CH4 hydrates are not stable above 200 m, glacial CH4 hydrates are actually stable at land surface, even as subsurface temperatures rise beneath the glacial blanket (Fig. 10a). During past glacial cycles, it is probable that methane hydrate formed and dissipated within the Canadian and Fennoscandian Shields (Stotler et al., 2010b). As methane hydrates are stable to depths of 2 km (e.g. Ershov and Yakushev, 1992), well beneath permafrost ice, and considering increased stability beneath glaciers (Fig. 10b), methane hydrate formation provides a possible explanation for freezing-evolved fluids well beneath the base of modern or paleo-permafrost, unrelated to glacial recharge. Considering methane hydrate formation also acts to reduce permeability (e.g. Schlu¨ter et al., 1998; Kleinberg et al., 2003), widespread CH4 hydrates formation prior to or during glacial advance would decrease the possibility for glacial meltwater to recharge deep into the subsurface, even if subsurface permafrost ice melted. However, methane is not ubiquitous at all crystalline shield sites; for instance at Ha¨stholmen (where freezing-evolved fluids are observed to 400 m), very little CH4 has been observed; thus at this location CH4 hydrate formation would have been limited to non-existent. The contribution of CH4 hydrates is similarly limited at Olkiluoto, because CH4 contents are very low at depths with glacial signals whereas high concentrations are observed several hundreds of meters deeper. Regardless, clearly understanding transient

gas hydrate stability fields could be a crucial component to glacial–interglacial cycling effects on groundwater recharge (in both shield and sedimentary units with high gas concentrations). Unfortunately, glacial recharge models have yet to consider the influence of methane hydrate. Methods utilized to obtain information from the deep subsurface can affect fluid flow in and around sample locations. Mine dewatering has affected hydraulic gradients and geochemical samples at sites across the Canadian Shield (e.g. Frape et al., 1984; McNutt et al., 1990; Douglas et al., 2000; Stotler et al., 2009). Significant changes in chemistry observed at the Con mine, even between the initial annual samples (collected in the 1980s), were partially attributed to strong downward gradients created during excavation and mine dewatering that resulted in significant fluid mixing in the rock mass (Frape et al., 1984; Douglas et al., 2000). The Con mine was in operation for the longest period of time prior to sampling and is also the site where freezing-evolved fluids were the deepest. However some sites were sampled through surface boreholes (East Bull Lake and Eye-Dashwa), or samples were collected prior to substantial changes to the subsurface flow field (Diavik). At Diavik, mixing was not an issue, and the effects of in situ freezing were evident in the chemistry (previously described). The other two sites were non-permafrost locations, with low d18O fluids found in relatively shallow freshwaters (Fig. 9b). These fluids could be relics of either glacial intrusion or the result of an in situ permafrost freeze-out process. Interpretation of fluid freezing mechanisms in the Canadian and Fennoscandian Shields must also take into account fluid density and solute transport. For instance, glacial meltwater recharge scenarios are complicated by the density and viscosity difference between the recharging and existing fluids, if glacial signals indicate infiltration to great depths. Dilute glacial fluids would have to displace the dense saline or brine fluids commonly observed in Shield rocks. Once displaced, a discharge location for the dense fluid already in the system must be found. Such density differences are currently accounted for in some glacial recharge models (Lemieux, 2006; Lemieux et al., 2008a,b,c). On the other hand, if permafrost formation were the only source of freezing-evolved fluids, downward transport of cryogenically concentrated fluids would be required, although several lines of evidence indicate this would not be the case. In crystalline rocks, groundwater transport is dominated by fracture flow and diffusion within the rock mass. Geochemical samples from within permafrost in crystalline rock indicate fluids are composed of a dilute, slightly oxidizing solution (Blowes and Logsdon, 1998; Stotler et al., 2009, 2011). Because salinities increase with increasing depth, downward diffusion of freezing-affected waters would not be significant. Rather, concentration gradients favor upward diffusion of solutes, towards the dilute fluid within the permafrost. Thermal gradients also favor upward fluid transport from deeper, warmer fluid, towards the colder permafrost (Cary and Mayland, 1972; Gray and Granger, 1986; Qiu et al., 1988; Perfect et al., 1991). Although geochemical effects of the glacial/interglacial cycle are found at considerable depths within the Canadian

Freezing processes and evolution of Shield groundwater

Shield, additional insight is needed to determine if glacial recharge, an in situ process, or anthropogenic influences such as mine dewatering are responsible for the geochemical trends observed to depths of 1100 m at individual sites. It is probable that no single freezing mechanism accounts for the freezing-affected portion of fluid evolution at every site. 4. CONCLUSIONS The glacial–interglacial cycle affects geochemical conditions in the deep subsurface. In this study, the available geochemical data from across the Canadian and Fennoscandian Shields were examined to determine residual geochemical effects from the glacial–interglacial cycle. Effects of glacial meltwater recharge and surficial freeze-out, previously considered by others, and in situ freeze-out effects due to ice and/or methane hydrate formation were examined. In instances where evidence of freezing was observed in fresh to saline fluids, the data were generally not sufficient to differentiate between mixed, intruded glacial meltwaters or residual waters resulting from an in situ freezing process (permafrost or methane hydrate formation). At some Canadian Shield sites, d18O values indicated that freezing-affected fresh, brackish, and saline fluids were found to depths of 1300 m. The geochemical characteristics associated with freezing processes were not observed at all sites. Freezing experiments and groundwater investigations in areas of permafrost provide guidelines for determining whether fluid evolution at a site has been affected by freezing. On a plot comparing Na/Cl and Br/Cl ratios, freezing of Na-dominated fluids results in only increasing Na/Cl ratios, but freezing of Ca-dominated fluids results in increases of both Na/Cl and Br/Cl ratios, contrary to hypothesized seawater freezing trends, and evaporative trends, with decreasing Na/Cl and increasing Br/Cl ratios. Fluid freezing results in an initial enrichment of d2H and d18O in ice. As freezing progresses, 2H and 18O continue to enrich in the ice until both ice and residual fluid are significantly depleted relative to the initial fluid. This fluid-ice mixture evolves along a line with a slope slightly less than the GMWL. During freezing, solutes concentrate in residual fluids, resulting in a linear relationship between increasing chloride and decreasing d18O on a semi-log plot. Mixing, as at the Con mine, results in a curved line on the same plot. It was clear that Canadian and Fennoscandian Shield brines did not evolve as a result of evaporation or the freeze-out process (surficial or in situ); brine evolution is more significant than fluid origin. Physical and geochemical data do not support the hypothesis that shield brines formed cryogenically in glacial marginal troughs as outlined by Starinsky and Katz (2003). Hydraulic gradients, subglacial pressures, and rock permeability data indicate it would be impossible to move a hypothetical brine from a marginal trough on the continental margin to the center of North America during the recent glacial cycles. Modern-day cryogenic brine formation in proglacial lakes is also not common in polar regions or near the margin of continental-scale ice sheets in Antarctica, Greenland, and North America. Geochemical evidence, including ionic relation-

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ships and isotopic (d2H, d18O, 87Sr/86Sr, d37Cl and d81Br) values indicate shield brines did not form as a result of surficial cryogenic freezing of seawater during Plio-Pleistocene glaciation. Geochemical freezing models also demonstrate that an in situ freeze-out origin for the brines is not possible, as subsurface temperatures in areas of permafrost are not currently low enough, and were not low enough in the past, to create such concentrated solutions. Modern subsurface temperatures in permafrost areas are also not low enough for the freeze-out process to have influenced any pre-existing concentrated shield brines. Methane was present at many of the study sites. It was suggested that formation and dissipation of paleo-methane hydrate at these locations could cause the observed dilute fluids, Na/Cl–Br/Cl, d2H–d18O, and Cl–d18O relationships observed at depths greater than paleo-permafrost, although evidence for geochemical modification of gas geochemistry was inconclusive. Methane hydrates could form during glacial advances due to increased pressures, reducing subsurface permeability, thereby reducing the potential for sub-glacial recharge. At this time, there is not enough evidence to support either methane hydrate formation or sub-glacial recharge as the main process influencing the observed geochemical trends. This research indicates further consideration of deep fluids across the Canadian Shield is necessary to understand geochemical evolution to support the safety case for safe radioactive waste disposal. Although shield brines were not formed through a freezing process, some of the more dilute fluids (<50 g L 1) have been affected. The possibility for deep glacial meltwater intrusion into the Canadian Shield versus an in situ freeze-out mechanism (either permafrost or methane hydrate formation) needs to be better constrained by collecting geochemical data from relatively undisturbed locations and analyzing recharge age and geochemical indications of freezing, and carefully accounting for mixing caused by anthropogenic disturbances. Pressure and salinity measurements from within hydrothermal taliks will provide an understanding of flow direction within taliks, and the relative importance of taliks to deep-flow-system evolution beneath the permafrost. Hydraulic glacial recharge modeling should incorporate talik formation and the potential for methane hydrate formation. A better understanding of the subsurface thermal history at individual sites, talik–permafrost interaction, and glacial rechargedischarge will further constrain processes involved in the evolution of fresh to moderately saline waters in the Canadian and Fennoscandian Shields. Two additional processes, in situ permafrost and methane hydrate formation, are introduced that create similar geochemical signatures as subglacial recharge, although at most locations it was not possible to distinguish between the processes. These processes should be considered in interpretations of freezing-evolved fluids and of paleo-groundwater simulations. ACKNOWLEDGMENTS Partial funding for this research was provided by a grant from the Natural Sciences and Engineering Research Council of Canada to the second author. Additional funding was provided by Ontario

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Power Generation of Canada in cooperation with SKB (Sweden), POSIVA (Finland) and the Geological Survey of Finland. The long-term support and access to the depths of the earth from the mining companies of Canada and Scandinavia was vital to the contents of this paper. Special permission granted by L. Smith at DDMI was especially appreciated. A review by A. Blyth and editorial comments from M. Adkins-Heljeson improved this manuscript.

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