Chemistry and mineralogy of Precambrian paleosols at Elliot Lake, Ontario, Canada

Precambrian Research, 12 (1980) 349--373

349

©Elsevier Scientific Publishing Company, Amsterdam -- Printed in The Netherlands

CHEMISTRY AND MINERALOGY OF PRECAMBRIAN PALEOSOLS AT ELLIOT LAKE, ONTARIO, CANADA

ALISON L. GAY* and D.E. GRANDSTAFF

Department of Geology, Temple University, Philadelphia, PA 19122 (U.S.A.) (Received August 10, 1979; revision accepted October 30, 1979)

ABSTRACT Gay, A.L. and Grandstaff, D.E., 1980. Chemistry and mineralogy of Precambrian paleosols at Elliot Lake, Ontario, Canada. Precambrian Res., 12: 349--373. Two profiles from the paleosol underlying the Huronian Supergroup at Elliot Lake, Ontario were studied to determine conditions present at the time of their formation. (1) One paleosol profile (> 10.5 m thick), developed on greenstone, is exposed within the Denison Mine. It is tentatively classified as an intrazonal gley soil formed in an area of poor drainage and reducing groundwater conditions. It is characterized by extensive loss of Fe, Mn, Cu, Ni, and Mo, as well as Na, Ca, and Mg. An increase in K is attributed to diagenetic K metasomatism. Paleosol structures indicate clay translocation forming an illuvial Bg horizon and suggests a fluctuating water table to depths of 6m. (2) The second paleosol profile (> 5.5 m thick), developed on granite, is exposed near the Pronto Mine. This profile exhibits iron oxidation and enrichment. This indicates that free oxygen was present in the mid-Precambrian atmosphere, although a precise determination o f p O 2 is not possible from evidence from the paleosols. This paleosol exhibits many characteristics of modern spodosols. The soils probably formed in a temperate humid climate during tectonically stable conditions. The difference in oxidation and behavior of iron in these soils is inferred to result from differences in paleotopography and drainage. The preponderance of gley paleosols found underlying the Huronian Supergroup might result either from lower mid-Precambrian pO~ or from preferential preservation. Evidence suggests that by 2.3 Ga/ago, soil forming processes were capable of producing soils similar to Recent types.

INTRODUCTION

The nature of the Precambrian earth surface environment has long been a subject of speculation and investigation. Speculation on mid- and late-Precambrian surface environments is constrained somewhat by geological evidence from Archean and Proterozoic sedimentary rocks. Precambrian banded iron formations, red beds, bedded cherts, and uraniferous conglomerates have all been studied intensely for evidence concerning conditions under which they formed. *Present address: Gulf Science and Technology, 11111 Wilcrest Drive, Houston, TX 77099, U.S.A.

350

Mature soils may also reflect conditions present at their formation. Studies of various Precambrian paleosols have been undertaken by Sharp (1940), Rankama (1955), Pienaar (1963), Roscoe (1969), Kalliokoski (1975), Fryer (1977), and others. Rankama (1955), Pienaar {1963), and Roscoe (1969) suggested that soil formation occurred under reducing conditions, possibly indicating an anoxic Precambrian atmosphere. Fryer (1977) suggested that soil formation occurred under alkaline conditions and suggested that Precambrian weathering may have been quite different from modern weathering. However, Sharp (1940) and Kalliokoski {1975) suggested that soil forming conditions were very similar to present conditions. In part the variety of conclusions may be related to the age of paleosol investigated. Kalliokoski and Sharp investigated later Precambrian paleosols (less than ca. 2 Ga age) whereas Rankama, Pienaar, Fryer, and Roscoe investigated older paleosols (greater than ca. 2 Ga age). This might suggest some evolution of soil forming processes. It is unfortunate that many of these earlier studies are incomplete or not sufficiently detailed. Many previous studies focused only on mineralogical or chemical analyses of a paleosol, considered only variations of a single element such as iron, or variations in the ferrous/ferric ratio, or examined only a limited number of samples. Study of both mineralogical and chemical variations in a paleosol is necessary to understand the environment under which it was formed. This paper reports the results of a mineralogical and chemical study of two early Aphebian (early Proterozoic) paleosols from Elliot Lake, Ontario, Canada. Results of this study suggest that by ca. 2.3 Ga/B.P. soil forming processes were capable of producing soils which resemble modern types. REGIONAL GEOLOGY

The Blind River-Elliot Lake region is located approx 30 km north of Lake Huron, half-way between Sudbury and Sault Ste. Marie, and 35 km northeast of Blind River, Ontario (Fig. 1). The geology of the area has been extensively studied, largely because of the presence of economically important Huronian uraniferous conglomerates. Results of various studies have been summarized by DeiTy (1960), Pienaar (1963), Roscoe (1969), Robertson (1978), and others. Rocks of the Blind River-Elliot Lake region include Archean granites and greenstones of the Whiskey Lake Greenstone Belt, overlain unconformably by Huronian sedimentary rocks, and intruded by Keweenawan lamprophyre and diabase (Fig. 2). The rocks have been folded into the broad westward-plunging Quirke Syncline and Chiblow Anticline. The Archean Whiskey Lake Greenstone Belt is an irregular west-northwesttrending belt approx 64 km in length. The belt is primarily comprised of basic and acidic volcanics, greywacke, conglomerate, iron formation, and various intrusives. The rocks underwent greenschist facies metamorphism and were later intruded by Archean granites and quartz diorites which are commonly

351

- ~.~

Contact

Fault

0

I0

20

30

Km ~

HURONIAN

ARCHEAN

~'Hlot Lake

Fig. 1. Locality map showing the contact between Arehean an Huronian rocks. Locations of the two paleosol profiles studied are given by designations Denison and Pronto. Locations of other paleosol profiles studied by Pienaar (P) and Dryer (F) are also shown. Roscoe (1969) has also studied a paleosol which crops out within the Pronto Mine. Major ore zones are outlined by dashed lines.

referred to as 'Algoman granite'. K/Ar measurements on the Arehean granites have yielded ages of ca. 2.4--2.7 Ga (Roscoe, 1969). The Huronian Supergroup consists of suites of subarkose, greywacke, argillite, limestone, and conglomerate. It has been subdivided into four groups, separated from one another by disconformities or non-conformities (Fig. 2). Maximum total stratigraphie thickness of the Huronian Supergroup is approx 8.5 km (Robertson, 1978). The Matinenda Formation of the Elliot Lake Group contains the economically important uraniferous conglomerates. These conglomerates generally lie stratigraphically within 100 m of the Archean/Huronian contact. There has been considerable debate concerning their origin. Although Davidson (1960, 1965), Derry (1960), Dimroth and Kimberley (1976) and others have suggested that the deposits are epigenetic, most mineralogical, geochemical, and stratigraphic evidence supports a detrital-placer hypothesis (Roscoe, 1969; Grandstaff; 1974b; Theis, 1978; Robert.son, 1978; and others). The Gowganda, Bruce, and Ramsay Lake Formations contain conglomeratic greywackes which have been interpreted by some geologists as tillites (Piennar, 1963; Roscoe, 1969; Young, 1970; Robertson, 1978). The Gowganda Formation has been studied extensively, and appears to contain the best evidence for a mid-Precambrian glaciation in North America. Sedimentary rocks of the Gowganda Formation have been dated at about 2.29 Ga (Fairbairn et al., 1969). Huronian and adjacent rocks throughout the region are intruded by sill-like bodies and dikes of gabbro, lamprophyre, and granophyre. Although these

352 E L L I O T LAKE STRAT[G}DkpHY

Nipissing Diabase

2. 15 Ga

Van S c h m u s

(1965)

2.29 GO

F a i r b a i r n et al.

2.4 - 2.7 Ga

Roscoe

HURONIAN SUPERGROUP Cobalt

(]roup

Bar River G o r d o n Lake Lorraine Gowqanda Qulrke

(1969)

Lake G r o u p

Serpent Espanola Bruce Hough Lake Group Mississagi Pecors Ramsay Lake Elliot Lake G r o u p McKim Matinenda Copper Cliff Thessalon L i v i n g s t o n e Creek ARCHEAN Paleosols Granite Greei~stonc

[1969)

Fig. 2. Simplified stratigraphic column and radiometric ages of Precambrian rocks in the Elliot Lake region (after Pienaar, 1963; and Roscoe, 1969).

rocks represent at least three different periods of intrusion, the name 'Nipissing diabase' has been applied generally to include all of them (Roscoe, 1969). These rocks have been dated at 2.15 Ga by whole rock Rb/Sr isochron (Van Schmus, 1965), and 2.1 Ga by K/Ar method on biotite (Roscoe, 1969). The contact between the Archean granite and greenstone and the Huronian conglomerate and arkose is marked by a paleosol (pedolith after usage o f Yaalon, 1971) ranging in thickness from a few cm to 20 m. The wide range o f paleosol thickness may result from localized soil erosion (Roscoe, 1969). The age of formation of these paleosols is probably 2.3--2.4 Ga. bracketed by the age of intrusion of the underlying parental granites and the age of the overlying Huronian Gowganda Formation. SAMPLES

Oriented samples were collected from two sections of paleosol (Fig. 1). One paleosol, developed on greenstone was sampled from a recently driven tunnel within the Consolidated Denison Mine. The second paleosol, developed on granite, was sampled from west of the Pronto Mine and near the discovery locality. In both cases the upper contact of the paleosol was with the Huronian Matinenda Formation. Sample depths were measured from the contact of the paleosol with the overlying Huronian sediments and perpendicular to the dip of the contact.

353 A N A L Y T I C A L METHODS

W~ole rock chemical analyses for major elements were performed by the "rapid" method of Shapiro (1967) with the following modifications or additions: silica was determined by molybdenum yellow technique, and ferrous iron was determined by titration with potassium dichromate (Maxwell, 1968). Ni and Mn were determined by atomic absorption spectroscopy after dissolution of the rock in HF-H2SO4--HNO3 solutions. Mo and Cu were analyzed colorimetrically using thiocyanate and 2,2'-biquinoline (Sandell, 1959). After pretreatment with HC1 to remove carbonates, reduced carbon was determined at Gollob Laboratories using a LECO TM carbon analyzer. Electron probe micro-analyses were made using an automated ARL-EMX electron microprobe. The measurements were made at 15 kV accelerating voltage with a 1--2 pm diameter beam. Glass and natural silicate standards were used. Data were reduced by the correction scheme of Bence and Albee (1968) and Albee and Ray {1970). P ET R OLOGY OF THE P A R E N T ROCKS

The Denison greenstone is a fine-grained, dark-green, massive to slightly foliated, meta-igneous rock which has undergone greenschist facies metamorphism. Major modal minerals are actinolite, chlorite, quartz, and sphene. Chlorite and actinolite both average Mg/(Mg+Fe) = 0.6. Minor or accessory minerals include epidote, microcline, rutile, leucoxene, hematite, biotite, muscovite, and pyrite. Relict amphibole phenocrysts, still showing characteristic cleavage angles and monoclinic crystal forms, have been replaced by epidote, quartz, sphene, and chlorite. The groundmass consists of masses of actinolite and chlorite and has a schistose texture. Fine-grained spheroids of sphene intergrown with hematite appear throughout the groundmass. These may represent replacement of ilmenite. Quartz grains rarely exceed a few millimeters in diameter. Most quartz grains have highly sutured boundaries with undulose extinction indicative of metamorphism. Some granular, recrystallized quartz is also present. The composition of the unweathered parent greenstone is somewhat unusual in that it contains < 2% total Na:O+K20+CaO (Table I). This results from the near absence of feldspar in the greenstone. The parent rock from which the Pronto paleosol was formed is a medium grained, leucocratic, sub-aluminous, pink alkali granite having an average grain size of 5 mm. The major modal mineralogy is quartz, microcline, and plagioclase (partially altered to sericite), with minor biotite, largely altered to chlorite even in the freshest samples. Accessory minerals include apatite and zircon. Microcline displaying albite-pericline twinning is the dominant feldspar; oligoclase is subordinate. Quartz grains have straight sutures and only slightly undulose extinction, indicating very low-grade metamorphism. The concentrations of A1203 and total iron are lower and the concentration of H20 ÷ higher

354 TABLE I Chemical analyses of samples of

SiO 2 (wt. %) TiO 2 A1203 FeO Fe203 MgO CaO Na20 K~O P2Os MnO H~O" H20 +

Mo (ppm) Ni

Depth (m) 0

0.3

0.6

63.94 0.93 21.03 1.87 2.60 0.96 0.07 0.45 5.92 0.27 0.022 0.11 2.25

71.72 0.20 13.75 1.44 2.30 1.34 0.07 0.43 5.02 0.39 0.012 0.33 3.24

73.06 0.27 12,49 1.72 1.85 1.85 0.16 0.30 4.59 0.39 0.015 0,15 2.47

72.42 0.42 12.94 3.16 1.78 1.17 0.17 0.27 4.38 0.10 0.019 0.27 2.72

69.70 0.61 12.75 5.97 0.34 1.58 0.19 0.35 3.16 0.41 0.048 0.35 4.11

67.78 0.86 12.60 6.49 2.20 1.52 0.16 0.26 3.17 0.42 0.070 0.18 4.12

3

4.5

100.42

100.24

99.32

99.82

99.57

99.83

4.6 32

5.0 35

3.6 32

9.1 29

12.2 22

11.0 32

38

28

35

53

35

17

16

65

153

230

2500

--

--

Cr

56

Cu

33

C (red)*

2

140

-1200

*Reduced ozganic carbon.

T A B L E II Chemical analyses of samples from northwest of the Pronto Mine Depth (m) 0

0.5

2.0

3.5

5.5

7.0

9.0

66.92 15.11 0.54 0.61 0.57 6.37 0.29 0.25 5.01 0.013 0.21 0.29 4.60 i00.78

77.50 10.93 0.37 0.39 0.28 2.48 0.17 0.13 4.54 0.004 0.28 0.13 2.34 99.54

72.20 14.54 0.21 0.66 0.29 1.30 0.20 0.15 7.26 0.005 0.40 0.15 2.22 99.59

72.76 13.61 0.21 0.90 0.43 1.02 0.19 0.18 7.70 0.006 0.31 0.12 2.76 100.20

73.66 12.70 0.09 0.64 0.72 0.99 2.92 0.48 5.72 0.005 0.39 0.14 1.86 100.32

72.53 12.04 0.13 0.31 0.48 0.03 4.36 0.26 5.42 0.011 0.33 0.14 3.82 99.86

73.76 11.54 0.14 0.42 0.57 0.21 4.18 0.31 5.43 0.013 0.63 0.11 3.10 100.41

F e 3 + / F e 2+

10.1

8.0

4.0

2.1

1.2

0.1

0.3

Ni ( p p m ) Cu

49 67

26 6.8

25 5.6

S i O 2 (wt. %) A1203 TiO 2 MgO FeO Fe203 Na20 CaO K~O MnO P205

H20H20 +

28 9.2

22 --

25 --

25 --

355

6

8.5

60.19 0.58 12.97 10.46 4.39 2.61 0.17 0.29 3.19 0.44 0.078 0.23 3.98

73.30 0.96 12.25 3.38 0.22 1.42 0.49 0.31 3.70 0.32 0.039 0.38 3.20

10.5

14.5

20

30

55.90 1.04 14.33 13.06 2.92 3.87 0.83 0.37 2.90 0.28 0.147 0.20 4.36

51.64 2.25 9.77 13.23 4.48 13.68 0.53 0.12 0.12 0.48 0.245 0.10 3.22

51.39 2.01 10.24 12.51 3.30 12.84 1.07 0.67 0.73 0.41 0.222 0.21 4.68

62.55 1.27 8.30 9.47 2.54 11.25 1.15 0.32 0.44 0.44 0.16 0.36 2.54

99.87

99.48

100.12

100.47

100.21

12.0 48 38

18.4 32 45

17.2 52 133

- -

1

1

0

- -

21.0 331 132

27.0 630 560

- -

- -

100.79 25.8 297 408 160

than the average alkali granite of Nockolds (1954) (Table II). The higher content of water probably reflects partial alteration of plagioclase to sericite and biotite to chlorite. PALEOSOLS

The contact between the Huronian Matinenda Formation and the underlying paleosol and Archean parent rocks could be easily recognized in the field by changes in color and texture of the rock, and by the absence of sedimentary structures and large pebbles within the paleosoh The absence of sedimentary (alluvial) structures in the paleosol suggests that the two profiles examined are residual deposits formed by in situ weathering and alteration of the parent rocks. Denison paleosol Figure 3 shows variations in Denison sample mineralogy with depth below the Archean/Huronian contact. The vast difference in mineralogy between samples above and below 10.5 m depth marks the base of the paleoweathering profile (paleosol). Thus the paleosol had a thickness of 10.5 m. This is a minimum thickness since some surface layers of the soil were undoubtedly removed by erosion. Lithostatic pressure may also have caused some post-formation compaction of the paleosol.

356

o.6I~ SRLI

1

Senclte

4.5

Quartz

,I

o

=

Actinolite ,

2O

40

Fe-

e

O

o%

6O

8O

I00

Fig. 3. Modal mineralogy of Denison Mine samples. Minor minerals include orthoclase (Or), epidote (E), sphene (Sph), biotite (B), and leucoxene (L). Other accessory minerals (stippled area) include rutile, pyrite, magnetite, and minor leucoxene.

During formation of the Precambian soil the minerals of the parent rock, actinolite, sphene, and epidote were completely altered or destroyed and replaced by clay minerals. These clay minerals were themselves recrystallized and altered to form fine-grained sericite. The clay content in the upper-most paleosol sample (as determined from the content of sericite) is > 50%. This is higher than in many modern soils. Quartz in the paleosol is a mixture of quartz residual from the greenstone and interstitial quartz cement. Some quartz grains in the paleosol are much larger than those observed in the parent greenstone. This suggests formation of quartz overgrowths. During formation of the paleosol magnesium~hlorite was altered to iron chlorite. Figure 4 shows the Mg/(Mg+Fe) ratio in chlorite from Denison samples as determined by electron microprobe analyses. In the parent greenstone, magnesian chlorite [Mg/(Mg+Fe) -~- 0.6] occurs within the groundmass and as large flakes. In the paleosol, magnesian chlorite is replaced by a fine-grained iron chlorite (Mg/(Mg+Fe) ~-- 0.35). Thus the chlorite which occurs within the paleosol is not residual but is a weathering product. Fryer (1977) suggested that iron chlorite would only occur below the paleowater table and postulated that its occurrence reflects a high paleowater table in the soil profile which he studied. However, Recent weathering of well-crystallized igneous or metamorphic magnesian chlorite c o m m o n l y results, initially, in formation of 0

o

~

3

oQ

~QO

6

8.5

o

0QD

2O 30

o13

I

I

I

0.5

I

•

"I

0.6

Mg /(Mcj -* Fe)

Fig. 4. Mg/(Mg+Fe) in chlorite from Denison samples as a function of depth as determined by electron microprobe analyses.

357

a fine-grained, poorly crystalline, iron-enriched "clay chlorite" (Loughnan, 1969). This is the t y p e of alteration seen in the Denison profile. This t y p e of alteration is c o m m o n under modern weathering conditions and is probably not indicative of water table position. In addition, some palosol structures suggest water movement to depths of ca. 6 m. Figure 5 shows the structural details of a polished section of paleosol. These structures are not inherited. We have found no textures like these in samples of the parent greenstone, nor do they appear to be metamorphic textures. The polishedsection appears to show individual angular to columnar ped units (soft crumbs), such as the large light colored triangular unit near the center of the section, up to ca. 1 cm in length. The soil structures are composed primarily of sericite, derived from clays, chlorite, recrystallized quartz, and quartz cement. Formation of such soil aggregates is usually aided by flocculation of clays and periodic dehydration (Birkeland, 1974; Soil Survey Staff, 1975). The texture is broken in places by cracks which may be dehydration or shrinkage cracks. The polished section was cut so that the top of the section is approximately parallel with the Archean/Huronian contact and with the Matinenda bedding. Therefore, many of these cracks were originally approximately vertical. These cracks appear to be filled with translocated and oriented clay skins which are n o w generally darkly colored. These clays appear to have been translocated (moved) from overlying soil horizons (an A or E horizon) and oriented by movement of soil waters. Oriented clay skins and M

Fig. 5. Polished section from 6 m below the Huronian/paleosol contact in Denison Mine. The polished section has been oriented so that the top of the polished section and the scale

are approximately parallel with Matinenda bedding. The scale is in cm. The polished section shows individual soil aggregate (peds), comprised dominantly of sericite, having a blocky to columnar structure: dehydration or shrinkage cracks (originally nearly vertical), and darkly colored illuvialclay films.

358 clay films appear in hand specimens, polished specimens, and thin sections of samples at depths from 1 to 6 m and suggest the formation of an illuvial horizon. Formation of illuvial features requires moving water and suggests that the permanent low-water table depth was no higher than 6 m. Although formation of illuvial features requires moving water, they apparently cannot form where water percolates through the soil at all seasons (Birkeland, 1974). Illuvial horizons are characteristic of lowlands where softs become thoroughly or partially dry at some season (Soil Survey Staff, 1975). Production of illuvial horizons probably requires more than 10 ~ years (Yaalon, 1971; Birkeland, 1974). Therefore land surfaces must have been relatively stable with low erosion rates. Mixing of soil by frost heaving or by shrinking and swelling must be slow or absent to permit these structures to form. This suggests that at the time of soil formation the climate, while seasonal, was not glacial or periglacial. Samples from the upper meter of the Denison profile did not contain clay films or other structures suggesting clay iUuvation. Although these structures can be obliterated b y diagenesis, their absence may indicate that remnants of an A or E soil horizon overlie the illuvial B horizon. Illuvial structures were not present in samples below ca. 6 m depth. This suggests that a C horizon underlies the B horizon at depths from 6 to 10.5 m. Samples from the Denison profile were taken only in one location; therefore, it is not known whether these structures and horizons are laterally continuous or h o w representative these horizon depths might be. Iron chlorite is c o m m o n (approx 15 modal %) in the lowest parts of the paleosol. However, it becomes much less c o m m o n in the upper meter (Fig. 3). This suggests that chlorite in the upper part of the soil profile was being destroyed by weathering. Iron chlorite is unstable in soils (Jackson, 1955} and would be expected to lose its hydroxide interlayers and break down at pH less than 6 under leaching conditions (Borchardt, 1977}. The pattern of persistence of iron chlorite suggests that the soil waters had pH greater than 6 at depth, decreasing near the surface to less than 6, resulting in partial breakd o w n of iron chlorite. Chemical analyses of the Denison samples are given in Table I and are displayed in Fig. 6. M o d e m softs contain organic matter. Near-contact samples from the Denison paleosol were analyzed for reduced organic carbon. Results are shown in Table I. Because these samples were taken underground from a recently driven shaft, they have probably undergone only slight recent organic contamination. The amount of organic carbon in the t o p 2 m of soil ranges from 0.014 to 0.25%. This amount is generally less than that in m o d e m soils, b u t is similar in abundance to organic carbon in some Archean shales (Moore and Welch, 1977). Organic c o m p o u n d s in soil are quite reactive and may easily be destroyed by oxidation or bacterial metabolism. Therefore, these values may be less than original amounts. Also, organic material tends to be concentrated at the t o p of the soil profile in O and A soil horizons; although high concentrations may also occur in illuvial B horizons. Because upper portions of the soil were probably eroded,

359 % ,0

% ,5

20

00.I

% 0.2

0.,0

%

0.5

;K20t /

L0

.

.

50 .

60

70

.

0.3

2

KzO 4

(%)

~i 6

&

0

o

s c°/°>,o

0.3

is o

o;s

Lo

,.s

s

It

I00

Ni,Cu,Cr (ppm) 200 300 400

I

,

!

I

500

600

I

I

0.6

0.6

tll'~l~ "E :3 "

3

CA)

FeO

,o':~\'~~___

Mo

~:~.4.5

,,.o~

' ~ - - ~ ~

....

,c,,,, Mo (ppm)

Fig. 6. Chemical analyses of Denison samples plotted as a function of sample depth below the contact. The dashed line at 14.5 m depth indicates the top of the parent rock as determined from mineralogical variations (Fig. 3).

360 layers of soil containing the highest content of organic material may have been removed. Although vascular plants had probably not evolved during the Precambrian, Fischer (1965), and Campbell and Golubic (in press) have proposed that algae or other microorganisms may have grown on or in subaerial soil profiles. Although the possibility of later organic contamination cannot n o w be excluded, the small amounts of organic carbon detected appear consistent with the presence of some organisms on or in the soil. Figure 6 shows chemical variations against depth. Many of the elements show major changes in concentration at or slightly above the 10.5 m sample, the base of the paleoweathering profile as determined from mineralogical changes. These changes in concentration can be ascribed to the effects of Precambrian weathering and subsequent diagenesis. CaO, MgO, MnO, TiO2, Ni, Cr, Cu, Mo, P2Os, FeO, Fe203, and FeO T (total iron as FeO) decrease from the parent rock to the paleosol, K20 increases, and Na20 shows no definite trend with depth. It is commonly assumed that Al is retained in the soil during weathering because it is relatively insoluble at normal pH values and much of it is incorporated in clay minerals (Birkeland, 1974). The concentration of Al generally increases during weathering because it is retained in the soil, while other components are removed by leaching. In order to assess more quantitatively the amount and nature of chemical weathering, we have constructed concentration ratio diagrams (Fig. 7) based on the assumption that aluminum is conserved within the soil profile during weathering. Other studies have occasionally assumed that Ti or Fe are conserved during weathering. However, the decrease in concentration of these elements in the Denison paleosol suggests that they were lost. Naturally, some aluminum is also lost during weathering, although its solubility and mobility are ordinarily low. Therefore, losses calculated are maxima. Figure 7 suggests that about 95% of CaO and MgO initially present in the rock were lost relative to alumina, losses similar to many Recent weathering profiles. The data also indicate ca. 90% loss of total iron and ca. 50% loss of Na. The amount of loss of major components relative to constant A1 indicated in Fig. 7 is: Ca ~-- Mg ~ Mn > Ti > Fe ~ P ~ Na > Si. While other elements are lost, potassium increases by approx. 600% relative to A1. The virtually complete removal of Ca and Mg in the Denison paleosol suggests that leaching was fairly intense. Under non-leaching (arid) conditions Ca and Mg are commonly retained in soil profiles as incrustations of calcite, gypsum, or other minerals. Although these incrustations and their Ca and Mg may be removed by diagenesis, desert incrustations tend to persist under changed environmental conditions or to leave textural evidence of their former presence (Yaalon, 1971). For example, Kalliokoski (1975) recognized caliche layers and textures indicating an arid climate from a late Precambrian paleosol in Michigan. Lack of obvious textures indicating caliche or arid conditions in the Denison paleosol suggests that loss of Ca and Mg is largely primary, occuring during Precambrian weathering, and indicates moderate to intense leaching conditions occuring under a wet or humid climate. Naturally some Ca

361 i

'

'

' ' ' ' ' l

,

CoO MgO

,

,

,

,

,,,

i

i

,

,

FeT TiO2 P20s

i , , ]

KtO

o

o.s ~

,

I

~o I

s

-r 4.5 0_ 6 o 8.5 I0.5 R

~

J

Loss i

,

0.03

i

,

i

J l J l

0.1

,

i

i

i

= i l l

0.3 Concentration Ratio

i

,

,

=

,

, , ,

3

I0

Fig. 7. Concentration ratio diagram for Denison samples showing gains and losses o f components relative to AI203. The concentration ratio (CR) is given by the equation:

c a = (M weathered/M parent) / (A120$weathered/A1203 parent ) where M weathered is the concentration of the oxide or element in a paleosol sample and M parent is the concentration in the parent rock. Components which plot on the left side of the diagram are lost; those that plot to the right side are gained with respect to Al203.

and Mg may have been lost during diagenesis. For example, formation of quartz cement, such as that n o w present in the samples, often involves replacement of pre-existing carbonate cement. Diagenetic alteration of clay minerals may also involve loss of Ca. In modern soils the behavior of silica can also be used to gauge the intensity of leaching. Under low to moderate leaching conditions, silica is retained in the soil profile in clay minerals, while under intense leaching conditions silica is removed in formation of lateritic or bauxitic soils. However, in these ancient paleosols the content of silica has probably undergone major diagenetic modification due to recrystallization of original clay minerals in formation of sericite and particularly in formation of quartz cement. Therefore, the present silica content of these paleosols is probably greater than during their formation and does not indicate the intensity of leaching conditions at the time of soil formation. The loss of Ti is unusual. Ti is generally retained during weathering and increases in concentration in soil. However, most Ti in the Denison parent rock is contained in sphene, which was destroyed by weathering. This liberated Ti was probably in the more mobile Ti(OH)4 species {Loughnan, 1969), which was then removed from the soil profile by ground water. Weathering usually results in changes in oxidation state of elements such as iron and manganese. The ferric/ferrous ratio in Recent soils generally increases due to oxidation during weathering. Rankama (1955} observed a decrease in the ferric: ferrous ratio of Precambrian paleosols from Suodenniemi, Finland, and concluded that anoxic conditions were present during their for-

362 mation. As shown in Fig. 6 the ferric: ferrous ratio in the Denison parent rock is approx 0.25. The ratio decreases to an average of 0.15 in the lower part of the paleosol (between 10.5 and 3 m below the contact), and then increases to as much as 1.7 at the contact. These variations suggest iron reduction at depth and oxidation toward the surface. Unfortunately iron oxidation states may change drastically during diagenesis. Diagenesis usually results in reduction of ferric to ferrous iron, with little change in the total iron content (Garrels and Mackenzie, 1971; Veizer, 1973). Therefore, the ferric: ferrous ratios of the samples are probably less than initial values. However, diagenetic changes in the ferric: ferrous ratio may be unpredictable unless the diagenetic history is well known. Therefore, conclusions drawn from these ratios alone may be suspect. Because diagenetic alterations influence the oxidation state b u t do not greatly alter bulk iron content, it may be more reasonable to rely on bulk changes in iron content to indicate oxidation or reduction during weathering. Table I and Fig. 6 show that total iron decreases from ca. 15% in the parent rock to 3--4% in the soil. If conditions are oxidizing during soil formation, iron liberated by dissolution of primary minerals is oxidized to insoluble Fe 3+ and precipitated in the soil profile as amorphous Fe(OH)3 or some other ferric iron mineral. Under reducing conditions iron is reduced to the more soluble Fe 2+, which may be removed by groundwater. Therefore, loss of iron from the paleosol suggests that during Precambrian weathering soil waters were anoxic or reducing and did not allow formation of insoluble Fe(OH)3 or other ferric phases. Modern soils which show iron reduction and loss are called gley soils {Soil Survey Staff, 1975). Gley soils generally also exhibit color mottling due to local variations in iron and manganese oxidation states. However, samples of the paleosol are a fairly uniform greenish or yellowish-green color. Because diagenesis alters oxidation states, color mottling based on these oxidation state variations would not be expected to survive (Yaalon, 1971). Cu, Ni, Cr, Mo, and Mn have all been lost from the paleosol. Their behavior parallels that of Fe. In many cases trace elements are co-precipitated with or adsorbed on iron-hydroxides, organic matter, or clay minerals, and are retained in the soil profile (Aubert and Pinta, 1977). Loss of these elements from the Denison paleosol suggests that organic matter (Table I) and ferric hydroxides were present only in small amounts or largely absent from the soil. The lack of ferric hydroxides suggested by these data is consistent with reducing conditions. The K20 content in the Denison profile increases from 0.45 in the parent rock to 6% at the contact (Table I, Fig. 6) and increases b y ca. 700% relative to A1 (Fig. 7). In contrast, modern softs show K depletion, the extent of which depends on the extent of illite formation. Recent weathering o f basic rocks under moderate leaching conditions generally results in formation of clay minerals such as montmorillonite, vermiculite, or halloysite, whereas intense leaching conditions may produce gibbsite or boehmite. It is unlikely that iUite could have formed to a major extent during weathering of the Denison greenstone due to a low initial content of K. However, clay minerals which may have originally been present in the soil have been totally replaced b y sericite.

363

Diagenetic alteration of original clay minerals to sericite would have resulted in the uptake of large amounts of potassium from groundwaters accompanied by smaller losses of Ca, Na, and Si (Garrels and Mackenzie, 1971). This diagenetic metasomatic alteration probably occured at low temperature. The Huronian Supergroup is virtually unmetamorphosed regionally (Roscoe, 1969). Although metamorphism does occur locally around the Nipissing Diabase intrusives, there is no evidence of local or contact metamorphism in the samples, nor does diabase crop out along the Consolidated Denison Mine tunnel from which the samples were taken. In summary: The Denison paleosol most closely approximates modern gley soils formed under anoxic or reducing groundwater conditions. The loss of iron suggests that groundwaters were generally anoxic, and variations in the ferric/ferrous ratio suggest that conditions were anoxic or reducing at depth but were at least partially oxidizing toward the surface. Soil structures suggest the presence of remnants of an A horizon in the top meter of paleosol, an illuvial B horizon from 1 to 6 m, and a C horizon from 6 to 10.5 m. Formation of illuvial structures and oriented clay films suggests that water percolated through the soil horizon for part of the year and suggests that the permanent low water table was at depths of ca. 6 m. The distribution of iron chlorite suggests slightly acidic groundwater near the surface, with neutral or alkaline conditions at depth. At the present soils having these characteristics form in tectonically and erosionally stable lowlands under temperate climate.

Pronto paleosol Figure 8 shows the sequence of mineralogical variations with depth in samples from the Pronto locality. Mineralogical and chemical variations indicate that the minimum thickness of the Pronto paleosol is approx. 5.5 m. Although samples of granite are slightly altered, even in fresh samples taken from far below the contact, the base of the paleosol can be recognized mineralogically by the rapid breakdown of plagioclase and alteration in rock texture from granitic (hypidimorphic-granular) to seriatic. Figure 8 shows that microcline altered much less extensively than did the oligoclase. In Recent weathering, plagioclase generally weathers faster than does potassium feldspar. The extent of alteration in the Pronto samples suggests that relative rates of weathering of these two minerals were the same in mid-Precambrian.

0

20

4()

°/o

60

80

I00

Fig. 8. Modal mineralogy of Pronto samples. Minerals include pyrite (P), magnetite (M), chlorite (Chl), and plagioclase (plag). Minor minerals indicated by the stippled area include leucoxene, rutile, apatite, and zircon.

364 Quartz grains in the paleosol are smaller than those in the parent granite. Aggregates of small quartz grains are c o m m o n in the paleosol and these aggregates form ca. 70% of quartz in the uppermost sample. This suggests that most of the quartz in the upper paleosol samples are not grains remnant from the parent rock, but were formed diagenetically. Relatively few opaque or heavy mineral grains occur in samples of the paleosol. Leucoxene, magnetite, and some rutile are scattered throughout, and pyrite comprises approx 2% of the uppermost paleosol sample. Biotite is almost completely replaced by chlorite in samples of parent granite. Some chlorite also occurs in the lowermost sample of paleosol, but not in higher samples. Hydrothermal replacement of biotite by chlorite is common and may occur soon after emplacement of the host rock. However, it cannot be unequivocally determined when chlorite replacement of biotite t o o k place in the Pronto samples. If replacement occurred prior to formation of the paleosol, the disappearance of chlorite from the paleosol may be due to acidic soilwater conditions as previously discussed. Modal analyses of thin sections and normative calculations from chemical analyses suggest that the content of clay minerals (now altered to sericite) is over 40%. If most of the quartz in the paleosol samples is diagenetic, the original clay content of the soil must have been higher, perhaps as much as 50 to 55%. Samples of Pronto paleosol do not show any diagnostic soil horizons although some soil structures such as soil peds are preserved. The absence of horizons in mineral soil can be a result of short periods of soil formation, or formation on steep, actively eroding slopes (Soil Survey Staff, 1975). Recrystallization could also have obliterated characteristic soil textures. Chemical analyses of Pronto samples are given in Table II and results are shown in Fig. 9. P, Ca, Na, and Fe 2+ decrease in concentration from the parent rock to the paleosol while Fe 3+, A1, K, Ti, Si, Mg, total iron, and the ferric/ ferrous ratio increase. The trace elements Cu, Mn, and Ni also increase near the surface. To assess further chemical weathering a constant ratio diagram was constructed under the assumption that A1 was conserved during weathering (Fig. 10). Figure 10 shows that ca. 95% of Na and 50% of Ca and P were lost relative to A1. K and Si are approximately constant relative to A1 while Mg may increase slightly. Ti increases by ca. 300% and total iron increases by as much as 700% relative to A1. The sequence of loss is approximately: Na > P ~ C a > Si ~ K ~ Al. The large decrease in Na content suggests that weatheringwas moderate to intense. Retention of K in the paleosol probably results from both formation of illite during weatheringand diagenetic metasomatismin replacement of clay minerals by sericite. The retention of silica in the paleosol probably results in part from diagenetic formation of quartz grains and cement. The content of total iron increases from < 1% in the parent granite to < 6% at the top of the paleosol, and the ferric: ferrous ratio increases. These suggest that iron was oxidized and retained in the soil profile. This suggests that

365

0

2

4

%

6

% 14

I0

8

18

o.s 2 --~ 3.5 ~ 5.~

i

7 9

60 0

2

4

6

8

o

;i02 SiO~[%) ppm 70

50

8()

I0O

150

Fig. 9. C h e m i c a l analyses o f P r o n t o samples p l o t t e d as a f u n c t i o n o f d e p t h b e l o w t h e contact.

groundwater conditions were oxidizing in contrast to reducing conditions in the Denison profile. The concentrations of Mn, Ni, and Cu increase at the t o p of the paleosol. Because these trace elements may be absorbed on ferric hydroxides, this increase is consistent with oxidizing conditions and formation of insoluble Fe(OH)3 or some other high surface area ferric phase. The constant ratio diagram (Fig. 10) indicates that Fe and Ti increase greatly relative to A1. Under alkaline conditions alumina becomes more soluble and might be lost relative to Fe and Ti. Therefore, these variations might be due to weathering under unusual alkaline conditions (pH > 9) such as those in modern playa lakes (Kalliokoski, 1975). However, high pH waters may also be highly undersaturated with respect to quartz and may rapidly and completely dissolve and remove quartz while leaving other minerals largely unaltered. Such alteration is n o t seen in the paleosol. Carbonates or other minerals may also precipitate from alkaline waters. Although these minerals may not persist, their presence usually results in formation of persistent soil textures not seen in this paleosol. Finally, a high pH environment would not be consistent with the distribution of chlorite in the paleosol, if chlorite were formed prior to soil formation. However, invocation of such extreme pH conditions may not be necessary. The increase in Fe and Ti relative to A1 could be due to translocation of these elements into the soil profile, b o t h in solution and organic

366

Na20

P205 CoO ,902 MgO

TiO,z

FaT

I

I\ aE:

Q03

QI

0.'3 Concentratioq~

I

3

I0

Ratio

Fig. 10. Concentration ratio diagram for Pronto samples showing gains and losses of components relative to Al~O3 (see caption of Fig. 7).

complexes or colloids, followed by precipitation or fixation in the soil profile (Birkeland, 1974). Craig and Loughnan {1964) and Goldich (1938) (see Fig. 47 and discussion in Loughnan, 1969) have observed increased contents of iron and titanium to A1 in Recent soil profiles developed on both basalt and gneiss under temperate conditions with normal pH. Recent soils which show good soil development, moderate leaching, absence of carbonate build-up at depth, and translocation of organic material, iron or other elements into the B horizon may be classified as spodosols (podsols) {Birkeland, 1974; Soil Survey Staff, 1975). Characteristics of the Pronto paleosol most closely approximate m o d e m spodosols. Elements are often translocated into the B horizon from overlying A or E horizons, but may also be translocated laterally by groundwater. In the Pronto paleosol, diagnostic horizon markers were not recognized. A or E horizons overlying a spodic B horizon contain lower concentrations of constituents translocated into the underlying horizon. In the Pronto paleosol Fe an Ti concentrations remain high up to the contact with the Huronian. Thus the B horizon appears to persist up to the contact, and O, A, or E horizons which may have overlain the B horizon appear to have been eroded. In summary: Characteristics of the Pronto paleosol most closely resemble modern spodosols. Increase in iron c o n t e n t and ferric : ferrous ratio suggests oxidizing conditions. Increases in Fe and Ti relative to A1 suggest translocation of these elements and fixation in a spodic B horizon. Extensive loss of Na suggests moderate weathering conditions. CONCLUSIONS

Although the two paleosols examined in the present study were formed within a few km of one another and at a b o u t the same time, they have quite

367 different characteristics. The Denison paleosol shows redution and loss of iron, and appears to have formed under reducing conditions, whereas the Pronto paleosol shows iron oxidation and enrichment and appears to have formed under oxidizing conditions. After their formation the softs were buried beneath sediments of the Matinenda Formation, which were deposited by a river or rivers flowing from the northwest (Roscoe, 1969). The paleosols studied probably formed along the floodplain. The occurrence of Precambrian paleosols exhibiting loss of iron and iron reduction (such as the Denison paleosol) has been interpreted as supporting the hypothesis of an anoxic or reducing early Precambrian atmosphere (Rankama, 1955; Roscoe, 1969). However, most recent speculation and geological evidence suggests that the mid-Precambrian atmosphere probably did contain some free oxygen, although the exact partial pressure is till a matter of intense debate (Holland, 1962, 1973, 1976; Berkner and Marshall, 1965, 1966; Brinkmann, 1969; Schidlowski, 1 9 7 6 ; D i m r o t h and Kimberley, 1976; Walker, 1977; Schopf, 1978; Towe, 1978; Dimroth and Lichtblau, 1978; and many others). Further, iron-reduced giey softs are n o w forming in swamps and in poorly drained areas along rivers such as the Mississippi and Indus. Therefore, it n o w appears more reasonable to interpret iron-reduced, gley paleosols, such as the Denison profile, not in terms of an anoxic Precambrian atmosphere, b u t in terms of local reducing or anoxic groundwater conditions caused by poor drainage. From drill core data it has previously been deduced that the Archean surface had considerable topographic relief prior to Huronian sedimentation (Hart et al., 1955; Derry, 1960; Roscoe, 1969). Huronian sediments were initially deposited in valleys between hills or ridges created by that relief. Relief may create lateral variations in drainage and lateral differences in soil t y p e (Birkeland, 1974; Bolt and Bruggenwert, 1976). Figure 11 shows a hypothetical toposequence which may conceptually represent conditions at the time of soil formation and burial by Matinenda sediments. The toposequence profile is characterized by upland areas having steeper slopes, a gently sloping pediment surface or piedmont plain, and a floodplain with levees and basins in the lowest parts. In areas of undulating topography, softs become progressively wetter on descending the slopes. This is known as a hydrologic sequence. At higher elevation drainage is better, groundwaters are oxidizing and most iron liberated by weathering is oxidized and retained in the soil. Depending on the nature of underlying strata, subsurface water from the highlands may move to great depths, or the water may reappear near the surface farther d o w n slope. Thus dissolved and colloidal material carried from sites in the highlands may influence softs at lower elevations. If iron, titanium, organic material or other constituents are translocated into the soil profile, a spodosol such as the Pronto paleosol may form under leaching conditions. At lower elevations drainage may be much poorer, or the soil profile may be flooded periodically b y river water or by seasonal storms, such as the monsoon in India and Bangladesh. Under flooded or poor drainage conditions, the soil waters become depleted

368

"%' ' ~ i ~

Fig. 11.Cross-section-toposequence showing generalized relationships between topography, fluvial sediments, and soil types. Fluvial sediments are shown by parallel curved lines. The area of possible development of gley soils is shown by the stippled area, spodsols by vertical dashes. Possible groundwater movement is shown by arrows.

of oxygen during various oxidation reactions and the soil waters b e c o m e anOxic or reducing. If the soil has a perennially water-saturated and reduced subsoil, while alternating reduction and oxidation take place near the surface, a gley soil may develop as in the Denison paleosol. Freyer (1977) concluded that rare earth distributions in an iron-reduced Elliot Lake paleosol indicated soil formation under alkaline conditions. He suggested that weathering must have been much different in the Precambrian. However, swampy reduced soils c o m m o n l y become alkaline at depth. Therefore, unusual conditions during the Precambrian may not be required to explain these data. Previous studies have suggested that reduced gley soils are fairly c o m m o n relative to oxidized soils in the mid-Precambrian, at least in the Elliot Lake region (Piennar, 1963; Roscoe, 1969; pers. comm., 1979; Fryer, 1977). We might expect that gley soils would be more common, that anoxic groundwater conditions necessary for their origin would be more easily created, if the oxygen content of the Precambrian atmosphere were lower than at present. Therefore, the relative preponderance of gley soils may be evidence of lower atmospheric oxygen pressure in the mid-Precambrian. This would be consistent with lower oxygen pressures which appear necessary for deposition of the overlying uraniferous placers of the Matinenda Formation and similar placers in the Witwatersrand. Although the survival of uraninite as a detrital mineral and formation of the uraniferous placers has been cited as evidence of a virtually anoxic Precambrian atmosphere (Ramdohr, 1958; Holland, 1962; Roscoe, 1969; Grandstaff, 1973, 1974a, in press) has calculated that the uraninite in the Elliot Lake and Witwatersrand deposits could have survived as a detrital mineral at oxygen pressures as high as ca. 0.01 PAL (0.002 atm).

369 However, the paleosols do n o t place strong constraints on pO2. The chemistry of the paleosol indicates that some free oxygen must have been present in the mid-Precambrian atmosphere. Amorphous Fe(OH)3, which probably was the initial ferric phase precipitated in the paleosol, can form at pO2 > 10 -13 atm (Klein and Bricker, 1977). However, oxygen pressures could have been much higher than the minimum value. However, it is also possible that the preponderance of gley softs is due to their preferential preservation, at least in the Elliot Lake area. Gley softs would have formed in lowlands near rivers, in a basin of depostion where they might be easily and completely preserved. More oxidized soils, like the Pronto Paleosol, would have formed on hills and ridges and areas of better drainage. However, these softs would be more likely to be removed by erosion, or further altered by subsequent pedogenetic processes. For example, it appears that the A horizon of the Pronto paleosol has been removed by erosion. Thus the predominance of gley soils might reflect preservational bias. Unfortunately, few early Phanerozoic or Precambrian paleosols have been investigated. In order to assess the possible importance of preservational bias, study of more early Phanerozoic and Precambrian paleosols is required. Some workers have inferred that the Huronian uraniferous conglomerates were deposited under cool or cold, glacial or periglacial conditions (e.g., Trow, 1977; Robertson, 1978). This inference is based in part on reports of tillites and tilloids of the Gowganda, Bruce, and Ramsay Lake Formations (Pienaar, 1963; Casshyap, 1969; Roscoe, 1969). Further, the persistence of uraninite as a detrital mineral and formation of the uraniferous conglomerate placers of the Matinenda Formation would be favored by glacial conditions {Grandstaff, 1973, in press; Simpson and Bowles, 1976). The tilloid stratigraphically closest to the Matinenda conglomerates comprises the Ramsay Lake Formation. Sedimentary rocks of the Huronian sequence have considerable lateral variation in stratigraphic thickness. In some areas the Ramsay Lake Formation may lie stratigraphically as much as 300 m above the. Matinenda conglomerates, while in others the Matinenda Formation may be truncated by tilloid deposits. It is likely that the maximum stratigraphic thicknesses represent areas of rapid deposition in subsiding zones whereas thinner areas represent areas o f lower sedimentation on topographic highs, or areas of erosion. Although sedimentation rates in molasse basins may exceed 100 c m / 1 0 0 0 yrs (Fischer, 1969), the average sedimentation rate on most Phanerozoic basins has been less than ca. 33 c m / 1 0 0 0 yrs. Thus the 300 m section between the uraniferous conglomerates and the tilloids of the Ramsay Lake Formation probably represents an interval of a b o u t 1 Ma. Obviously much climatic change can occur in such a period of time. Therefore, the tilloids of the Gowganda, Bruce, and Ramsay Lake Formations may not be a good guide to climate at the time of Matinenda deposition. Paleosols too may provide an indication of climate at the time of their formation. Most paleosols investigated in the Elliot Lake region (Pienaar, 1963; Roscoe, 1969; Fryer, 1977; this study) appear to have been intensely weathered

370

and to have originally had a high clay content. This clay has been diagenetically replaced by sericite. The Pronto and Denison paleosols examined in this study must have had initial clay contents of approximately 40--50% or more. The clay content of soils varies with climate under which they form. Jenny {1935, 1941) attempted to find quantitative relationships between climatic factors and the clay content of soil. He suggested that the rates of clay production vary linearly with moisture and exponentially with temperature. Thus soils formed under arid and cold climates generally contain little clay, and are mostly comprised of rock fragments. Soils formed in tropical, h o t and rainy climates contain large amounts of clay or aluminum and iron hydroxides or oxyhydroxides. The relationships were developed by Jenny before many of the complexities of Pleistocene stratigraphy became known; therefore, although the generalized trends are probably correct, the quantitative relationships are probably not valid (Birkeland, 1974). Although it is not possible to determine absolute values, based qualitatively on the relationships of Jenny, the high clay content of the Denison and Pronto paleosols suggests a temperate climate with moderate to high rainfall. The illuvial structures suggest that rainfall was seasonal. Because the paleosols were probably formed shortly before Matinenda deposition, they may be more valid than the tilloids as indicators of climate at the time of deposition of the uraniferous conglomerates. The rate of Precambrian weathering and soil-forming processes is largely unknown. Because plants were probably much less abundant in subaerial soils than at present, weathering may have been somewhat slower unless the scarcity of plants was compensated by an increase in atmospheric CO: pressure (Cawley et al., 1969}. Soils may only form in areas in which soil formation is faster than denudation. Variations in paleosol thickness at Elliot Lake have suggested that material was removed and that the area was undergoing some erosion (Roscoe, 1969}. Although the rate of denudation may have been slow due to low relief (Ahnert, 1970), the formation of thick soils in areas of at least slow denudation suggests that the soil forming process was not very much slower than at present. ACKNOWLEDGEMENTS

The authors are indebted to Denison and Rio Algom Mining corporations for access to their properties. We wish to thank R.J. Gunning, Chief geologist of Denison mine, for his help in sample collecting. We also wish to thank Drs. V. Ruzicka and J. Friel for their aid, and M. Shapiro for typing the manuscript. Support for this study was provided by Temple University.

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